Minor Alkaline Earth Element and Alkali Metal Behavior in Closed-Basin Lakes

Home , Anhydrite, Carnallite, Debye (crater), Fryxell (crater), Lake Hoare, McKnight Creek

Dissertation

Presented in Partial Fulfillment of the Requirements for the Degree Doctor of Philosophy in the Graduate School of The Ohio State University

Rebecca A. Witherow

Graduate Program in Geological Sciences

The Ohio State University

2009

Dissertation Committee:

W. Berry Lyons, Advisor

Anne Carey

Ozeas Costa

David Porinchu

William Green

This page intentionally left blank.

Abstract

Hydrologically closed basins in arid environments often pose ideal conditions for the development of inland saltwater lakes. In these regions, where the hydrology is balanced by dilute inflow waters and evaporation, these closed-basin lakes may progress through various degrees of evapoconcentration. This process may result in the precipitation of simple salts and the alteration of the relative chemical composition of a lake. Such “geochemical evolution” processes have been studied in detail particularly for the so-called major ions (Na, Ca, K, Mg, Cl,

SO4, HCO3+CO3) and Si, but little work has been done on ions of moderate concentration. Due to the evaporative nature of these systems, the alkali metals, Li and Rb, and alkaline earth elements,

Sr and Ba, may become highly concentrated and play a significant role in determining the geochemistry of a saline lake. Here, I present the first comprehensive study of the transport and fate of the minor alkali elements, Li, Rb, Sr, and Ba, in three distinct geographic settings illustrating three types of saline lakes as determined by their major anion abundance. Using lithium isotope analysis and mass balance calculations, I show that the minor elements in the

McMurdo Dry Valleys are transported to “chloride-type” lakes by a combination of precipitation, chemical weathering and salt dissolution. Sr and Ba show evidence of removal in the lakes, but

Li and Rb appear to be conservative to extremely high concentrations. This work discusses the potential removal mechanisms of these ions, particularly the formation of major and minor sulfate and carbonate minerals in the lakes. A major finding of this study is that brine type determines the removal mechanism of Sr and Ba, particularly if they are removed as sulfate or

ii carbonate minerals. This study provides a platform for future exploration of the role of these and other minor elements in inland brines and playas.

iii

Dedicated to my parents.

Acknowledgments

I wish to thank my adviser, W. Berry Lyons for his guidance, insight, and patience. Berry has been an amazing mentor and has taught me much about science and what it means to be a scientist. He has been supportive of my ideas, and he has given me direction when I’ve needed it. I am forever grateful for the opportunities he has given me. He has helped me achieve my goals of seeing the world, including Antarctica and the Arctic, and obtaining my Ph.D.. It has been a true honor working with him through the years.

My thanks go to my committee members, Anne Carey, Ozeas Costa, Dave Porinchu and

William Green for their scientific and stylistic comments and suggestions. Their unique perspectives have greatly enhanced this dissertation, and I am indebted to their commitment.

I also wish to thank Kathleen Welch, Joel Barker, Sarah Fortner, and Christopher Gardner for their friendship, field assistance, scientific knowledge, and editorial comments. Their advice support has made me a better scientist, and I will always have fond memories of our times together.

It has been a pleasure working with the past and present members of the McMurdo Dry

Valleys Long Term Ecological Research project. The collaboration within the group is inspirational, and I admire their commitment to understanding the McMurdo Dry Valleys ecosystem. I am especially grateful for the collection of lake samples by the Montana State Limno

Team and for the University of Colorado Stream Team for stream sample collection. I am indebted to the support staff in the McMurdo Dry Valleys especially Rae Spain and Sandra Liu.

Their hospitality at Lake Hoare camp is unparalleled, and they have made Antarctica feel like home.

This work could not have been completed without the assistance of several people. John

Olesik and Anthony Lutton from TERL at OSU were extremely helpful assisting in ICPMS analysis. Gideon Henderson and Andrew Mason at Oxford University were essential in lithium isotope analysis and interpretation. I am grateful for the generosity of William Last for donating lake samples to this project.

This journey through graduate school has been long and difficult, and I am sincerely thankful for the moral support of my family and friends. Thank you Shawn, Mom, and Dad for listening to me, letting me cry, and telling me I’m smart. Thanks Casey and Laura for listening to my stories. To all of my friends going through the same thing, thank you for making me feel like

I’m not alone.

This work was supported by NSF-ANT grant #0423595, a GSA Graduate Student

Research Grant, and an OSU Office of International Affairs Travel Grant.

Vita

2001 ...... B.A. Geology, Miami University

2005 ...... M.S. Geological Sciences, The Ohio State University

2002-present ...... Graduate Teaching and Research Associate, The Ohio State University

Publications

1. Harmon, R.S., Lyons, W.B., Long, D.T., Odgen, F.L., Mitasova, H, Gardner, C.B., Welch, K.A., Witherow, R.A. (2009) Geochemistry of four tropical montane watersheds, central Panama. 24(4), Applied Geochemistry.

2. Witherow, R.A. and Lyons, W.B. (2008) Mercury deposition in a polar desert ecosystem. Environmental Science and Technology. 42(13), 4710-4716.

3. Witherow, R.A., Lyons, W.B. (2008) The role of minor alkali elements in the evolution of closed-basin lakes. GSA Abstracts with Programs. 40(6).

4. Witherow, R.A., Henderson, G., Lyons, W.B. (2008) Lithium isotopes in a polar desert, McMurdo Dry Valleys, Antarctica. Goldschmidt Conference Abstracts. A1029.

5. Fortner, S.K., Lyons, W.B., Witherow, R.A., Welch, K.A., Olesik, J.W. (2007) Trace metal dynamics and transport in a polar glacier-dominated watershed: Taylor Valley, Antarctica. Goldschmidt Conference Abstracts 2007. A289.

6. Witherow, R.A., Lyons, W.B., Welch K.A. (2007) Geochemical controls on major and minor alkaline elements in closed-basin lakes, Taylor Valley, Antarctica. GSA Abstracts with Programs. 39(6).

7. Lyons, W.B., Harmon, R.S., Gardner, C., Welch, K.A., Witherow, R.A., Long, D.T. (2006) The geochemistry of headwater streams, central Panama: Influence of watershed rock type. GSA Abstracts with Programs. 38(7), 96.

vii

8. Witherow, R.A., Lyons, W.B., Bertler, N.A.N., Welch, K.A., Mayewski, P.A., Sneed, S.B., Nylen, T., Handley, M.J., Fountain, A. (2006) The aeolian flux of calcium, chloride and nitrate to the McMurdo Dry Valleys landscape: Evidence from snow pit analysis. Antarctic Science. 18(4), 497-505.

Field of Study

Major Field: Geological Sciences

viii

Table of Contents

Abstract ...... ii Acknowledgments ...... v Vita ...... vii List of Figures...... xi List of Tables ...... xv Introduction ...... xvi Chapter 1 Geochemical Dynamics of Minor Elements in Closed-Basin Lakes in Antarctica ...... 1 1.1 Abstract ...... 1 1.2 Introduction ...... 1 1.3 Site Description ...... 4 1.4 Methods ...... 6 1.5 Results ...... 10 1.5.1 Major ion concentrations ...... 10 1.5.2 Concentrations of minor elements in streams ...... 12 1.5.3 Minor element concentrations in lakes ...... 15 1.5.4 Minor elements in lake sediments ...... 21 1.6 Discussion ...... 32 1.6.1 Minor alkali metals and alkaline earths in MCM streams ...... 32 1.6.2 Lake chemistry ...... 40 1.6.3 Geochemical sinks ...... 48 1.7 Conclusions ...... 51 1.8 Acknowledgements ...... 52 Chapter 2 Lithium Isotopic Composition of the mcmurdo Dry Valleys Aquatic Systems ...... 53 2.1 Abstract ...... 53 2.2 Introduction ...... 54 2.3 Site Description ...... 56 2.4 Methods ...... 58 2.5 Results ...... 61 2.6 Discussion ...... 67 2.6.1 Glaciers and aeolian input ...... 67 2.6.2 Streams: Chemical weathering and salt dissolution ...... 69 2.6.3 Lakes and isotopic mixing ...... 71 2.7 Conclusions ...... 80 2.8 Acknowledgements ...... 80

ix Chapter 3 The Fate of Minor Alkali Elements in the Chemical Evolution of Closed-Basin Lakes 82 3.1 Abstract ...... 82 3.2 Introduction ...... 82 3.3 Site Description ...... 85 3.3.1 McMurdo Dry Valleys ...... 85 3.3.2 Great Basin ...... 88 3.4 Methods ...... 92 3.4.1 Sample collection ...... 92 3.4.2 Laboratory analysis ...... 93 3.4.3 Geochemical modeling ...... 94 3.5 Results ...... 95 3.1.1. Concentrations: Major elements ...... 95 3.5.1 Concentrations: Minor elements ...... 99 3.5.2 PHREEQ ...... 104 3.6 Discussion ...... 105 3.6.1 Inorganic mineral precipitation ...... 106 3.6.2 Replacement and incorporation into other minerals ...... 110 3.6.3 The potential importance of biogenic precipitation ...... 115 3.7 Conclusions ...... 116 3.8 Acknowledgements ...... 117 Chapter 4 Conclusions ...... 118 Appendix A: Map of the McMurdo Dry Valleys ...... 122 Appendix B: Detection Limits, Precision Data, and Blank Concentrations ...... 125 Appendix C: Mean Major and Minor Ion Concentrations of Streams ...... 128 Appendix D: Minor Ion Concentrations for the Lakes in This Study ...... 136 Appendix E: Map of the Great Basin ...... 141 Appendix F: Map of the Saskatchewan Lakes ...... 143 Appendix G: Major Ion Concentrations of the Lakes in This Study ...... 145 Appendix H: Saturation Indices for the Minerals and Lakes Discussed in Text ...... 150 References ...... 156

List of Figures

Figure 1.1 Major ions in the MCM streams and lakes (mol%)...... 11

Figure 1.2 Chloride and lithium concentrations in MCM streams. The trend line represents best first-order fit...... 13

Figure 1.3 Chloride and rubidium concentrations in MCM streams. The trend line represents the best first-order fit...... 14

Figure 1.4 Chloride and strontium concentrations in MCM streams. The trendline represents the best first-order fit...... 14

Figure 1.5 Chloride and barium concentrations in MCM streams. The trendline represents the best first-order fit...... 15

Figure 1.6 Longitudinal profile of the Onyx River. Closed symbols represent concentration in μg/L. Open symbols represent the minor ion concentration to chloride concentration ratio in g:mg...... 17

Figure 1.7 Longitudinal profile of Canada Stream. Closed symbols represent concentration in μg/L. Open symbols represent the minor ion concentration to chloride concentration ratio in g:mg...... 18

Figure 1.8 Lithium concentration by depth for the MCM lakes...... 22

Figure 1.9 Li:Cl values for the MCM lakes. Seawater from (PILSON, 1998)...... 22

Figure 1.10 Rubidium concentration by depth for the MCM lakes...... 23

Figure 1.11 Rb:Cl values for the MCM lakes. Seawater from (PILSON, 1998)...... 23

Figure 1.12 Strontium concentration by depth for the MCM lakes...... 24

Figure 1.13 Sr:Cl values for the MCM lakes. Seawater from (PILSON, 1998)...... 24

Figure 1.14 Sr:Ca values for the MCM lakes. Seawater from (PILSON, 1998) ...... 25

Figure 1.15 Sr:Mg values for the MCM lakes. Seawater from (PILSON, 1998) ...... 25

Figure 1.16 Barium concentration by depth for the MCM lakes...... 26

Figure 1.17 Sr:Cl values for the MCM lakes. Seawater from (PILSON, 1998)...... 26

xi Figure 1.18 Lithium concentration contained in the adsorbed (blue), sulfate (red), and carbonate (green) phases of MCM lake sediments. *Only adsorbed fraction analyzed. **Only adsorbed and sulfate fractions analyzed...... 30

Figure 1.19 Rubidium concentrations contained in the adsorbed (blue), sulfate (red), and carbonate (green) phases of MCM lake sediments. *Only adsorbed fraction analyzed. **Only adsorbed and sulfate fractions analyzed...... 30

Figure 1.20 Strontium concentrations contained in the adsorbed (blue), sulfate (red), and carbonate (green) phases of MCM lake sediments. *Only adsorbed fraction analyzed. **Only adsorbed and sulfate fractions analyzed...... 31

Figure 1.21 Barium concentrations contained in the adsorbed (blue), sulfate (red), and carbonate (green) phases of MCM lake sediments. *Only adsorbed fraction analyzed. **Only adsorbed and sulfate fractions analyzed...... 31

Figure 1.22 Magnesium and lithium concentration in MCM streams...... 35

Figure 1.23 K to Rb relationship in stream water...... 36

Figure 1.24 Cl versus K/Rb relationship in stream water...... 36

Figure 1.25 Sr to Ca relationship in MCM stream water...... 38

Figure 1.26 Potassium to barium concentrations in MCM stream water...... 39

Figure 1.27 Bicarbonate + carbonate concentration versus calcium and magnesium concentration for the MCM lakes and streams. Open symbols represent calcium concentrations, and solid symbols represent magnesium concentrations. Dashed line is a 1:1 ratio. HOR = Lake Hoare; FRX = Lake Fryxell; ELB = East Lake Bonney; WLB = West Lake Bonney; VAN = Lake Vanda; DJP = Don Juan Pond. Although the bicarbonate + carbonate concentration Don Juan Pond data are below the detection limit, it is plotted along the y-axis such that its calcium and magnesium concentrations can be compared to the streams and other lakes. ... 42

Figure 1.28. Chloride concentration versus calcium and magnesium concentration for the MCM lakes and streams. Open symbols represent calcium concentrations, and solid symbols represent magnesium concentrations. Dashed line is a 1:1 ratio. HOR = Lake Hoare; FRX = Lake Fryxell; ELB = East Lake Bonney; WLB = West Lake Bonney; VAN = Lake Vanda; DJP = Don Juan Pond...... 43

Figure 1.29 Li:Mg ratios in the MCM lakes. Note that Don Juan Pond is not included on this plot to better examine the variation in the other lake profiles. Seawater data (Pilson, 1998) ...... 46

Figure 2.1 Sodium concentration by depth for the lakes examined in this study...... 63

Figure 2.2 Lithium concentration versus δ7Li. Lake surface water samples are lighter colored. With the exception of Don Juan Pond and seawater, bottom water samples are darker Seawater data are from (PILSON, 1998; ROSNER et al., 2007)...... 65

xii Figure 2.3 Lithium to calcium mass ratios versus δ7Li. Lighter symbols represent lake surface water samples, darker symbols represent bottom water samples, with the exceptions of Don Juan Pond and Seawater. Seawater data from (PILSON, 1998; ROSNER et al., 2007)...... 65

Figure 2.4 Lithium to sodium mass ratios versus δ7Li values. Light colored symbols represent lake surface water samples, dark colored symbols represent bottom water samples, with the exceptions of Don Juan Pond and Seawater. Seawater data from (PILSON, 1998; ROSNER et al., 2007)...... 66

Figure 2.5 Lithium concentrations and isotopic values for other terminal lakes. The blue diamonds are from this study. The red squares are from (TOMASCAK et al., 2003). The green triangles are from (CHAN and EDMOND, 1988). Seawater data from (PILSON, 1998; ROSNER et al., 2007)...... 74

Figure 2.6 Seawater-glacier melt mixing model for both monimolimnia of Lake Bonney. Lake Bonney samples are green squares. Seawater is a blue asterisk. The average of 3 glacial snow samples is represented by a blue triangle. The tick marks on the linear regression line represent the relative amount of seawater in the hypothetical mixture...... 75

Figure 3.1 The geochemical evolution of closed-basin lakes as determined by major ion concentration of waters and minerals precipitated (HARDIE and EUGSTER, 1970)...... 84

Figure 3.2 The locations of the sites examined in this study...... 86

Figure 3.3 Piper diagram of the lakes in this study...... 97

Figure 3.4 Chloride and lithium concentrations in the lakes in this study. SK=Saskatchewan, GB=Great Basin, MCM=McMurdo Dry Valleys, SW=seawater (PILSON, 1998)...... 101

Figure 3.5 Chloride and rubidium concentrations in the lakes in this study. SK=Saskatchewan, GB=Great Basin, MCM=McMurdo Dry Valleys, SW=seawater (PILSON, 1998)...... 101

Figure 3.6 Chloride and barium concentrations in the lakes in this study. SK=Saskatchewan, GB=Great Basin, MCM=McMurdo Dry Valleys, SW=seawater (PILSON, 1998)...... 102

Figure 3.7 Chloride and strontium concentrations in the lakes in this study. SK=Saskatchewan, GB=Great Basin, MCM=McMurdo Dry Valleys, SW=seawater (PILSON, 1998)...... 102

Figure 3.8 Calcium and strontium concentrations in the lakes in this study. SK=Saskatchewan, GB=Great Basin, MCM=McMurdo Dry Valleys, SW=seawater (PILSON, 1998)...... 103

Figure 3.9 Carbonate mineral saturation indices for the MCM lakes...... 108

Figure 3.10 Sulfate mineral saturation indices for the MCM lakes...... 109

Figure 3.11 Mirabilite saturation indices for the MCM lakes...... 110

xiii Quadrant IV: calcite is supersaturated, and strontianite is undersaturated. B) Quadrant I: calcite and witherite are supersaturated. Quadrant II: witherite is supersaturated, and calcite is undersaturated. Quadrant III: witherite and calcite are undersaturated. Quadrant IV: calcite is supersaturated, and witherite is undersaturated. C) Quadrant I: gypsum and celestite are supersaturated...... 114

xiv

List of Tables

Table 1.1 Major sources and sinks of lithium, rubidium, strontium, and barium in the global ocean. 1(Stoffyn-Egli and Mackenzie, 1984); 2 (James and Palmer, 2000); 3(DeHairs et al., 1987); 4(DeHairs et al., 1980); 5(Edmond et al., 1979); 6 (Drever, 1988)Based on the weathering of igneous material and the formation of sedimentary material; 7(Elderfield and Gieskes, 1982); 8(Palmer and Edmond, 1989a); 97 (Elderfield and Schultz, 2006); 8(Broecker, 1971) ...... 4

Table 1.2 Reagents used in the sediment leaching experiments ...... 9

Table 1.3 Detection limits and precision analyses for sediment leaching experiments...... 9

Table 1.4 Lithium, rubidium, strontium, and barium concentrations in the adsorbed, sulfate, and carbonate phases in MCM lakes...... 28

Table 1.5 Average lithium, rubidium, and strontium concentrations of stream sediments analyzed by XRF...... 29

Table 2.1 Duplicate δ7Li Analysis...... 61

Table 2.2 Lithium, sodium, and calcium concentrations and δ7Li values for McMurdo Dry Valley glaciers, streams and lakes. Seawater concentration data from (PILSON, 1998) and isotope data from (ROSNER et al., 2007)...... 64

Table 2.3 Average lithium and sodium composition of stream sediment samples compared to stream water samples...... 66

Table 2.4 Lithium concentrations and isotopic values for other closed-basin lakes. Mono Lake δ7Li value is a mean of 10 samples and Li concentration is a mean of 6 samples. Walker Lake data are a mean of 2 samples...... 73

Table 3.1 Reactions, products and phases added to the Pitzer database in PHREEQ ...... 96

Introduction

Endorheic basins are located in semi-arid to arid regions on all continents, and their hydrologies are determined by the delicate balance of the input of water through springs, streams, and very limited precipitation, and the loss of water through evaporation or sublimation. The balance of these physical properties has a profound effect on the chemistry of desert lakes particularly when: 1) the outflow of water is restricted such that the lake is hydrologically closed; 2) evaporation exceeds inflow; and 3) the inflow is sufficient to provide a standing body of water (EUGSTER and HARDIE, 1978). Under these constraints, the geochemistries of closed-basin lakes are highly sensitive to natural and anthropogenic perturbations. For example, sediment cores retrieved from Pyramid Lake, Nevada have been used to identify regional droughts in the Great Basin since the Holocene (BENSON et al., 2002a). Over the past century, much of the Truckee River water has been diverted for agricultural use resulting in a

77% reduction in volume and a 160% increase in total dissolved solids (TDS) (GALAT et al., 1981).

This scenario is not unique to the Great Basin as, globally, waters feeding closed basin lakes are being diverted for human use (VITOUSEK et al., 1997).

The ultimate chemistry of a so-called closed-basin lake is determined at a very early stage by the weathering reactions between dilute inflow water and lithology and subsequent evapoconcentration of the lake (HARDIE and EUGSTER, 1970). The major evolutionary trend in brine formation is determined by the early precipitation of calcite (CaCO3) through evapoconcentration, it is at this first “geochemical divide” that the lake becomes either enriched

xvi or depleted in calcium relative to alkalinity (HCO3+CO3). As a lake becomes more evapoconcentrated, gypsum, mirabilite, and/or sepiolite precipitation relatively depletes or enriches Ca and Mg in the brine (EUGSTER and HARDIE, 1978). It is clear that calcium and magnesium play a critical role in the chemical evolution of closed-basin lakes. Yet, the behavior of minor alkaline earth elements and alkali metals in saline environments is poorly understood and predominantly limited to studies in marine systems.

The overall objective of my dissertation is to quantify the availability and mobility of alkaline earth elements and alkali metals in saline lake waters. My central hypothesis is that the behavior of minor alkali metals and alkaline earth elements is determined by the transport of these elements from dilute inflow waters to closed basin lakes, the degree of chemical evolution of the lake, the relative abundances of major ions, and removal mechanisms of minor cations. I have tested this hypothesis by examining precipitation, stream water, and lake water chemistries of the McMurdo Dry Valleys of Antarctica (MCM), the Great Basin of the western United States (GB), and southern

Saskatchewan, Canada (SK). Included in this dissertation are three manuscripts that discuss various aspects of the geochemistry of these sites.

Chapter 1 evaluates the concentrations of lithium, rubidium, strontium, and barium in the streams and lakes in the MCM. The aim of this chapter is to remark on these hypotheses:

 Hypothesis 1a: Like the major ions in MCM streams, the sources of minor alkali

metals and alkaline earth elements are the result of precipitation, chemical

weathering, and salt dissolution

 Hypothesis 1b: Biological uptake and adsorption onto clays are sinks for minor

alkalis.

xvii Chapter 2 examines the lithium isotopic ratios of glacial snow, stream, and lake water in the McMurdo Dry Valleys to address the following hypotheses:

 Hypothesis 2a: Lithium isotopic ratios will identify the source of lithium to the

MCM streams.

 Hypothesis 2b: The fresher surface waters of the lakes will be isotopically similar

to the streams reflecting the sources on the modern lithium inputs.

 Hypothesis 2c: The older hypersaline waters will have unique signatures as a

result of the climatic history of the lake.

In Chapter 3, I have examined the behavior of lithium, rubidium, strontium, and barium in three major anionic brine types. I have manipulated the major and minor ion concentration data from the lakes with the geochemical model PHREEQ to determine the saturation indices of major and minor minerals in the lakes. This chapter addresses the following hypotheses:

 Hypothesis 3a: Mineral saturation indices will reflect the degree of brine evolution

 Hypothesis 3b: Brines of different major anion type will display differences in

major and minor mineral saturation indices

 Hypotheses 3c: Examining the saturation indices of major and minor carbonate

and sulfate minerals will indicate potential sinks for alkaline earth elements and

alkali metals in closed systems.

xviii

Chapter 1 Geochemical Dynamics of Minor Elements in Closed-Basin Lakes in Antarctica

1.1 Abstract

The behavior of solutes in saline lakes has been of interest to geochemists for decades. Of particular curiosity is the so-called geochemical evolution of these lakes where dilute inflow waters evaporate to produce different chemistries. For the most part, these studies have focused on the “major” ions typically in highest concentration in the inflow waters: Na, K, Ca, Mg, Cl,

SO4, and HCO3+CO3. Little work has been conducted on elements of moderate or “minor” concentration in inflow waters. Here I have quantified the concentration of Li, Rb, Sr, Ba, and Cs in streams and lakes in the McMurdo Dry Valleys, Antarctica. Although Cs concentrations were all below the detection limit, these data add valuable information to previous research on the transportation and fate of solutes in the Dry Valleys. Li, Rb, Sr, and Ba come from a combination of marine salt dissolution and chemical weathering. Based on the relationships between these minor elements and major elements, Ba and Rb are being actively removed in streams, while Li and Sr are remaining in solution and are being conservatively transported to the closed-basin lakes. There is evidence to suggest that Li and Rb are being removed from Lake Fryxell, a brackish lake with an anoxic hypolimnion. Sr and Ba are not conservative in Lake Vanda.

1.2 Introduction

The primary influence on the initial composition of a lake is the bulk lithology of the catchment, dissolution of soluble minerals, leaching of adsorbed ions, and silicate weathering

(JONES, 1966). In closed-basin lakes, after solutes are delivered, evaporation causes the initial 1 composition to change, as the least soluble salts are precipitated, leaving behind brines that differ from inflow waters (HARDIE and EUGSTER, 1970). The chemical pathways leading to the ultimate composition of a closed-basin lake brine are determined at a very early state in the development of a closed-basin lake and are linked to the composition of dilute inflow waters (HARDIE and

EUGSTER, 1970). There are few studies that have identified minor elements in saline lakes: Br in the Dead Sea (5,000-7,000 mg/L), Sr in the Great Salt Lake (2,000 mg/L), PO4 in Searles Lake (900 mg/L), and B in Borax Lake (350-900 mg/L) (EUGSTER and HARDIE, 1978), yet, little research has been done on the sources and sinks of minor elements and their geochemical behavior in saline lakes. This study examines the sources and fate of minor alkaline elements and alkali metals in the McMurdo Dry Valley system.

The hydrology of McMurdo Dry Valleys (MCM) is relatively simple compared to watersheds in more temperate locations. The source of stream water is almost exclusively glacial melt (FOUNTAIN et al., 1999b). There are virtually no inputs from direct precipitation and few streams have tributaries. The streams flow directly into ice-covered endorheic lakes. The hydrologic budgets of these terminal lakes are dictated by the gain of liquid water by steam input and the loss of water by water freezing on the bottom of the lake ice covers and its subsequent sublimation. Thus the lakes vary in size with climatic variations, especially changes in temperatures in the austral summer (DORAN et al., 2002). During the Holocene, the lakes in the

McMurdo Dry Valleys varied greatly in their extent from much larger to much smaller than they are today (LYONS et al., 1998b; POREDA et al., 2004; SMITH and FRIEDMAN, 1993; WILSON, 1964)

Understanding the controls on the concentrations of minor alkali elements in saline lakes may provide important insight into the minor element mechanisms occurring in other natural waters. Like continental endorheic basins, in the ocean there is no outflow of water and evaporation plays a primary role in the hydrologic balance. The concentration of a solute can be 2 dependent on of biotic and abiotic processes and salinity (BROECKER and PENG, 1982). To establish whether the chemical balance of the oceans is in equilibrium or steady state, the sources and sinks of solutes must be evaluated. The inputs and outputs of major elements (e.g. Na, K,

Mg, Ca, Cl, and SO4) has been compiled in detail in earlier works (e.g. (BROECKER and PENG, 1982;

DREVER, 1988; PILSON, 1998). Despite the diverse mechanisms for removing these elements from the water column, the chemical mass balance between streams, hydrothermal waters, and oceans indicates that there may be a process removing cations that has not been accounted for

(MACKENZIE and GARRELS, 1966). It has been suggested that cation uptake by authigenic clays, termed “reverse weathering,” is responsible for this removal, but empirical evidence of this mechanism has been limited (GAC et al., 1977; GARRELS and MACKENZIE, 1967; VONDAMM and

EDMOND, 1984)

Crustal spreading centers in the oceans have a major control on the geochemical balance of alkali elements in the ocean (Table 1.1). Along with riverine inputs, high temperature

“hydrothermal” fluids entering the ocean at spreading centers are the primary inputs of Li, Rb,

Sr, and Ba. These mid-ocean ridges can not only be an important source of solutes, but they provide a constant supply of fresh basalt susceptible to chemical weathering and cation removal.

For example, hydrothermal fluids are an important source of lithium and cation exchange reactions on fresh basalt is a dominant removal mechanism (STOFFYN-EGLI and MACKENZIE,

1984). The investigations of closed-basin lake geochemical dynamics may aid in our overall understanding of cation removal in the oceans (VONDAMM and EDMOND, 1984). This research aims to discuss the potential sources and removal mechanisms, including reverse weathering, occurring in hydrologically closed-basin lakes perhaps analogous to the global ocean.

3 Sources Sinks

Riverine input Atmospheric cycling Li1 Hydrothermal Low temperature basalt alteration Diffusion from sediments Metamorphism

Riverine input Estuary mixing Rb2 Hydrothermal Low temperature basalt alteration

Riverine input8 Hydrothermal8 Carbonate formation8 Sr Alteration of volcanics6 Low temperature basalt alteration7 Carbonate recrystalization6

Riverine input5 Estuary mixing3 Ba3 Atmospheric dust3 Barite settling3 Hydrothermal5 Biotic production of barite4

1.3 Site Description

The McMurdo Dry Valleys (MCM) is the largest ice-free area in Antarctica

(approximately 4,800 km2) (Appendix A). It is a series of east-west trending valleys situated between the Transantarctic Mountains and the Ross Sea. The hydrology of the valleys is driven by summer glacial melt as there is not overland flow due to precipitation and no groundwater

(MCKNIGHT et al., 1999). The region is classified as a polar desert with mean annual temperatures of approximately -20°C and precipitation is less than 10 cm weq/year (DORAN et al., 2002;

FOUNTAIN et al., 1999b; WITHEROW et al., 2006). This precipitation usually falls as snow in the accumulation zones of the glaciers, and most of the snow that reaches the valley floor sublimates limiting its contribution to stream flow (CHINN, 1981).

4 The valley floor is a mosaic of alpine glaciers, saline soils, ephemeral streams, and perennially ice-covered lakes. The major rock types in the valleys consist of dolerites, granite, basalts, gneisses, schists, sandstones, and metasediments (CLARIDGE and CAMPBELL, 1977). Past glacial advances and retreats have reworked these lithologies into features such as drifts and moraines (Denton et al., 1971). The present climate has further altered the morphology producing polygons and desert pavements. Compared to glaciers in temperate regions, MCM glaciers are relatively clean with very little visible sediment on the surface and at depth. The primary source of solutes to the glacier surfaces are the aeolian deposition of dust from the valley floor and marine aerosols (FORTNER et al., 2005; LYONS et al., 2003; WITHEROW et al., 2006).

For approximately 6-10 weeks during the summer, meltwater flows from the terminus of the glaciers forming streams of various lengths and gradients (CONOVITZ et al., 1998). The flora in the streams is limited to algae and moss (MCKNIGHT et al., 1999). The primary source of solutes to the streams is through chemical weathering in the stream channel and hyporheic zone, the dissolution of salts in the stream channel, and inputs from solutes from glacial melt (FORTNER et al., in review; GOOSEFF et al., 2002; GREEN et al., 1989). Although the stream channels are surrounded by permafrost, active hyporheic zones contribute to the dissolved stream load

(MCKNIGHT et al., 1999). Beyond the wetted zones, there is an apparent lack of groundwater in

Taylor Valley (MCKNIGHT et al., 1999). This observation coupled with the absence of seasonal snow cover indicates that the dry valley stream input is the only source of solutes to the MCM lakes (CONOVITZ et al., 2006).

Much is known about the major ion geochemistry of the lakes in the MCM (GREEN et al.,

1988; GREEN et al., 1989; GREEN and LYONS, 2009; LYONS et al., 1998b). The differences in the geochemistries and salinities of these lakes are largely due to the differing water loss

(evaporation/sublimation) and drawdown histories linked to climatic variations (LYONS et al., 5 1998b; MATSUBAYA et al., 1979; POREDA et al., 2004). The Taylor Valley Lakes have been the primary focus of the National Science Foundation’s McMurdo Dry Valleys Long Term Ecological

Research project (MCM-LTER) (Appendix A). Lake Hoare is the freshest of the lakes with Cl concentrations ranging from 31.9 to 206 mg/L with increasing concentrations with depth. Lake

Fryxell is brackish and well stratified; its Cl concentration ranging from 386 to 3,480 mg/L. A submerged sill separates the two lobes of Lake Bonney, both of which are chemically stratified.

The Cl concentration in the western lobe of Lake Bonney ranges from 383 mg/L to 88,600 mg/L, and the concentration in the eastern lobe ranges from 355 mg/L to 188,000 mg/L. Lake Vanda and

Don Juan Pong are in Wright Valley to the north. The chloride concentration in Lake Vanda ranges from 259 mg/L to 70,900 mg/L with a distinct chemocline at 60 m below the ice surface.

Don Juan Pond is a shallow hypersaline brine whose salinity may daily and seasonally. The salinity has been shown to be as high as 40 wt% primarily as Ca and Cl (HARRIS and

CARTWRIGHT, 1981; TORII and YAMAGATA, 1981).

1.4 Methods

Following the protocols of the MCM-LTER, 149 stream samples were collected from 23 streams during the austral summers of 2005/2006 and 2006/2007. Depth profiles from 4 lakes were collected during the 2005/2006 summer. Water samples analyzed for cations were collected in NalgeneTM polyethelyne bottles that had been soaked with 10% HCl and rinsed 5 times with 18 mΩ water (DI). Water samples analyzed for anions, were collected in similar bottles that had been rinsed 3 times with DI. Stream samples were collected by submerging the bottle into the on- coming flow in the stream and rinsing it three times with sample water taking care to discard the rinse water downstream. A Niskin bottle was deployed to various depths in the lakes and retrieved water was used to fill pre-cleaned samples bottles. To minimize contamination

6 between lakes, separate Niskin bottles and filtering equipment were used for each lake. Samples were collected such that no head space remained in the bottle. All samples were filtered through

0.4 µm Nuclepore filters, and cation samples were acidified with HNO3 (0.1% for streams 1% for lakes) within 24 hours of sampling. Major anion and major cation samples were stored chilled at

4° C until they were analyzed at Crary Laboratory at McMurdo Station within 2 weeks of collection. Samples to be analyzed for minor elements were sent to The Ohio State University and stored in the dark at room temperature.

Samples were analyzed for F, Cl, SO4, Na, K, Mg, and Ca on a Dionex DX-120 Ion

Chromatograph (IC) using the techniques of (WELCH et al., 1996). A stock standard solution was created from several single-element standards. The multi-element stock standard was diluted to produce six calibration standards with a range of concentrations that covered the estimated concentrations of the samples. An independent multi-element standard was analyzed after creating a calibration curve to satisfy quality assurance and quality control. Several sample duplicates were analyzed for recovery comparisons and were used for precision analysis

(Appendix B). Clean sampling bottles filled with DI, i.e. “field blanks,” were transported to sampling sites, briefly opened, and then filtered and analyzed by the same techniques as with the samples themselves (Appendix B). This was done to determine if sample processing affected the contamination of the samples. As seen in Appendix B, the concentrations of ions in the field blanks were all less than 2% of the concentration in the sample.

Samples from the McMurdo Dry Valley lakes were analyzed for dissolved inorganic carbon (DIC). The samples were bottled with no head space and preserved with 0.5% v:v chloroform. The samples were stored chilled in the dark until analysis by either a Lira IRGA with an HP integrator or a Licor 6252 system in the Crary Lab (MCM-LTER, 2007). Because the lakes are stratified, DIC replicate measurements below the chemoclines can be compared to 7 illustrate the precision of the method (Appendix B). Streams were analyzed for alkalinity by titrating with HCl (STANDARD METHODS, 1999). Independent analyses conducted in 17 different laboratories, show that the reproducibility of samples up to 120 mg/L CO3 + HCO3 was within 5%

(STANDARD METHODS, 1999)

Stream and lake samples were analyzed for the minor elements, Li, Rb, Sr, Ba, and Cs, on an Elan Quadrupole ICPMS at the Trace Element Research Laboratory at The Ohio State

University. A multi-element stock standard of Li, Rb, Sr, Ba, and Cs was diluted to the appropriate range to create a six-point calibration curve. During stream sample analysis, a multi- element standard was analyzed every 5 to 10 samples to monitor instrument drift. High salinity lake samples were diluted and analyzed using an internal standard of a 100 μg/L solution of Be,

Co, Y, and La to monitor and correct for instrument drift. As a means of determining the precision of the analyses, several sample duplicates and check standards of a moderate concentration were examined for reproducibility (Appendix B). As with the major ions, several field blanks were analyzed to ensure my sampling methods did not introduce an unacceptable level of contamination (Appendix B). The detection limits for the ions analyzed by ICPMS were determined by multiplying the relative standard deviation of the lowest standard by 3 (Appendix

B). All samples had Cs concentrations below the detection limit of 0.05 µg/L, and therefore Cs will not be discussed in this paper.

Leaching experiments were performed to analyze for the concentrations of minor alkalis in lake particulate matter using the techniques of (FROELICH, 1980; KELLER et al., 2007; RUTTEN and DE LANGE, 2002; SCHENAU et al., 2001). Particulate matter samples were obtained from sediment traps in the lakes, and one sample was analyzed from the surface sediment of Lake

Fryxell. These samples were mixed in an effort to homogenize the grains before 3 20 g aliquots were weighed. These samples were reacted with various reagents (i.e. leaching solutions) 8 targeting specific fractions of the samples. These solutions were collected using qualitative cellulose filters (pore size = 20 µm). The first aliquot was rinsed with 30 mL of 2M MgCL2 (Table

1.2). The second aliquot was rinsed with 30 mL of 2M MgCL2 followed by 90 mL of 1 M NH4Cl

(to pH 7 with NH4OH). The final aliquot was rinsed with 30 mL of 2M MgCL2, 90 mL of 1 M

NH4Cl (to pH 7 with NH4OH), and finally 30 mL of 10% H3PO4. The leachates were then analyzed by Elan Quadrupole ICPMS at the Trace Element Research Laboratory at The Ohio

State University. To overcome matrix effects, standards were made using the various leaching reagents, and the reagents themselves were analyzed to ensure no contamination. All analyses were done using an internal standard of a 100 μg/L solution of Be, Co, Y, and La to monitor and correct for instrument drift. As a means of determining the precision of the analyses, several sample duplicates and check standards of a moderate concentration were examined for reproducibility (Table 1.3). The detection limits for the ions analyzed by ICPMS were determined by multiplying the relative standard deviation of the lowest standard by 3 (Table 1.3).

Target Fraction Reagent Volume (mL) Adsorbed1 1 M MgCl2 30 Sulfates2 2 M NH4Cl (to pH 7 with NH4OH) 90 Carbonates3 10% H3PO4 30

Table 1.2 Reagents used in the sediment leaching experiments. 1. (RUTTEN and DE LANGE, 2002), 2. (KELLER et al., 2007; SCHENAU et al., 2001), 3. (FROELICH, 1980)

Li (µg/L) Rb (µg/L) Sr (µg/L) Ba (µg/L) Detection limit 0.001 0.001 0.009 0.005 1 M MgCl2 DL DL DL DL 2 M NH4Cl (to pH 7 with NH4OH) 0.004 DL DL DL 10% H3PO4 0.006 0.002 DL DL Maximum duplicate variability 13% 6.7% 5.0% 14%

Table 1.3 Detection limits and precision analyses for sediment leaching experiments. 9 1.5 Results

1.5.1 Major ion concentrations

Streams in the MCM typically had Ca or Na + K as the major cation, and Mg was usually found in relatively low concentrations (Appendix A, Figure 1.1). In most of these streams, Cl and

HCO3 + CO3 accounted for most of the negatively charged ions. Some streams, particularly

Andersen Creek and Santa Fe Creek, which flow from or close to Blood Falls, a hypersaline discharge, had high proportions of SO4 (Appendix C) (LYONS et al., 2005).

The lakes in the dry valleys had different chemistries than the streams, and this was largely dependent on their age and stage of chemical evolution (Figure 1.1, Appendix C). Lake

Fryxell and Lake Hoare had high relative concentrations of Cl (71-90%). Lake Hoare had low but approximately equal amounts of SO4 (11-13%) and HCO3 + CO3, (11-17%), but Lake Fryxell had little SO4 (0.7-4.6%). In both Lake Fryxell and Lake Hoare, Na + K were the dominant cations, ranging from 68% to 88% of the total. Cl was the dominant anion in the more saline lakes, Don

Juan Pond, Lake Vanda, and Lake Bonney where Cl accounted for 90-100% of the total anions. In both lobes of Lake Bonney, Na was the cation in the highest concentration (58% to 81%), but at and below the chemocline in East Lake Bonney, Mg accounted for between 35 and 40% of cations.

Lake Vanda had Ca as the major cation at and below the chemocline (52% and 51%), but in the epilimnion, Na + K was more prevalent (49%). Ca in Don Juan Pond was almost the only cation found in the lake (85%). In all of the lakes, SO4 concentrations were depleted relative to inflow streams.

100 80 60 40 20 0 0 20 40 60 80 100 Ca Cl + F Lake Bonney Bonney Streams Don Juan Pond Lake Fryxell Fryxell Streams Lake Hoare Hoare Streams Lake Vanda Vanda Streams Seawater

Figure 1.1 Major ions in the MCM streams and lakes (mol%).

1.5.2 Concentrations of minor elements in streams

Lithium concentrations in the streams ranged from 0.09 μg/L to 5.01 μg/L (Appendix C).

The median concentration was 0.65 μg/L, and the mean concentration was 1.08 μg/L, which is similar to the world average concentration of 1.84 μg/L(Appendix C) (GAILLARDET et al., 2003).

Lithium had a strong correlation with chloride concentration in all of the streams (r2 = 0.77; n =

149) (Appendix A, Figure 1.2). Huey Creek had the highest average concentration of lithium

(3.36 μg/L) and the highest concentration of an individual sample (5.01 μg/L). Mariah Stream had the lowest Li concentration of an individual sample (0.09 μg/L). Like Huey Creek, Mariah

Stream is in the Lake Fryxell basin, but Mariah Stream is much shorter and has a lower slope than

Huey Creek.

The mean rubidium concentration of the stream samples was 0.83 μg/L (median = 0.68

μg/L) (Appendix C). This concentration is lower than the world average of 1.63 μg/L

(GAILLARDET et al., 2003). Rubidium concentrations in the streams ranged from below the detection limit of 0.07 μg/L to 4.53 μg/L. Two samples from Lyons Creek and Lawson Creek had concentrations below the detection limit, and the highest Rb concentration was detected in a sample from Santa Fe Creek. Of the streams that were sampled on multiple occasions, Priscu

Stream had the highest average Rb concentration (1.62 μg/L; n = 4) and Green Creek had the lowest (0.44 μg/L; n = 5). Unlike lithium, rubidium concentrations in stream waters do not show a strong correlation with chloride concentration (Figure 1.3).

The mean and median strontium concentrations of the MCM streams were 56.7 μg/L and

37.9 μg/L, respectively (Appendix C). This average Sr concentration of the MCM streams was only slightly lower than the world average of 60.0 μg/L (GAILLARDET et al., 2003). The maximum

12 Sr concentration value was in a Delta Stream sample (190 μg/L), and the minimum (5.68 μg/L) was from a Lawson Creek sample. Crescent Stream had the highest average concentration of Sr

(166 μg/L; n = 5), and Lawson Creek had the lowest (8.18 μg/L; n = 4). Strontium concentrations in streams showd a fair correlation with chloride concentrations (r2 = 0.61; n = 149) (Figure 1.4).

The mean concentration of barium in the streams was 1.63 μg/L, and the median concentration was 1.17 μg/L. The average concentration was significantly lower than the world average of 23 μg/L (GAILLARDET et al., 2003). A sample from Lost Seal Stream had the lowest barium concentration of the streams in this study (0.11 μg/L) (Appendix C). A Santa Fe Creek sample had the maximum concentration of barium of the MCM streams (15.8 μg/L). Of the streams with multiple samples, Santa Fe Creek had the highest average barium concentration

(5.01 μg/L; n = 5) and McKnight Creek had the lowest (0.66 μg/L; n = 4). Barium concentrations showed no correlation with chloride (Figure 1.5).

g/L)

μ Li ( Li 0.1

y = 0.25x0.75 R² = 0.77 0.01 0.1 1 10 100 Cl (mg/L)

Figure 1.2 Chloride and lithium concentrations in MCM streams. The trend line represents best first-order fit.

10 g/L)

μ 1 Rb ( Rb

y = 0.48x0.25 R² = 0.20 0.1 0.1 1 10 100 Cl (mg/L)

Figure 1.3 Chloride and rubidium concentrations in MCM streams. The trend line represents the best first-order fit.

1000

100

g/L)

μ Sr ( Sr 10

y = 5.08x + 20.57 R² = 0.61 1 0.1 1 10 100 Cl (mg/L)

Figure 1.4 Chloride and strontium concentrations in MCM streams. The trendline represents the best first-order fit.

100

g/L)

μ Ba ( Ba 1

y = 0.85e0.03x R² = 0.06 0.1 0.1 1 10 100 Cl (mg/L)

Figure 1.5 Chloride and barium concentrations in MCM streams. The trendline represents the best first-order fit.

In Wright Valley, the Li:Cl values in the Onyx River in Wright increased downstream, and Rb: Cl ratios increased downstream until the boulder pavement at 31.5 km downstream when the Rb:Cl values increased. In Canada Stream, Li:Cl and Cr:Cl values decreased from 0-437 m downstream, after which, they remained constant. In the Onyx River, Sr to Cl ratios increased slightly from approximately 0-15 km after which, they remain constant. In Canada Stream, Sr:Cl values remain fairly constant throughout the length of the stream. In both streams, Ba:Cl longitudinal profiles show no discernable pattern.

1.5.3 Minor element concentrations in lakes

Lithium concentrations for Lakes Hoare, Fryxell, and Bonney have been previously quantified and my data compared well to this earlier work (LYONS and WELCH, 1997). Lithium concentrations in the McMurdo Dry Valley lakes ranged from 2.47 μg/L to 2.65 x 105 μg/L

(Appendix D). The highest concentration was in Don Juan Pond and the lowest was from Lake

15 Hoare at 4 m. Lake Hoare Li concentrations ranged from 2.47 μg/L to 16.8 μg/L with the lowest concentration at 4 m and increased with depth (Figure 1.8). The Li concentration in Lake Fryxell was higher than that in Lake Hoare, ranging from 23.6 μg/L at 6 m and increasing until 99.1 μg/L.

The surface water of both lobes of Lake Bonney had relatively low Li concentrations: 29.3 μg/L in the east lobe and 22.4 μg/L in the west lobe. At depth in both lobes, the brines had concentrations of to 8,290 μg/L Li in the east lobe and 4,400 μg/L Li in the west lobe. Lake Vanda also displayed a large Li concentration gradient from 22.4 μg/L in the surface water to 22,000 μg/L at 70 m. Li:Cl values were highest in Don Juan Pond followed by Lake Vanda. The Li:Cl values in all of the lakes were fairly constant with depth with the exception of Lake Fryxell , which decreased with depth (Figure 1.9).

Longitudinal profiles were collected from Canada Stream (12/16/2006) and the Onyx

River (12/18/2006) (Figure 1.6Figure 1.7). Flow in Canada Stream (31.3 to 41.1 L/s) was average

(mean = 35 L/s) during the sampling period. The Onyx River’s flow (272 to 312 L/s) was approximately average (mean = 234 L/s) for the sampling period as well. The Onyx River was only flowing in the main channel, and no tributaries were feeding the river. In both streams, Li,

Rb, and Sr concentrations generally increased downstream, but in the Onyx River, Rb concentrations decreased between 20.8 km and Lake Vanda. In both streams, Ba showed no relationship between distance and concentration.

In the lakes and pond, Rb concentrations ranged from 2.11 μg/L to 294 μg/L (Appendix

D). Lake Hoare had the lowest concentration of Rb at 4 m, with concentrations increasing with depth (Figure 1.10). All of the Lake Fryxell samples had higher Rb concentrations than Lake

Hoare and ranged from 6.38 μg/L at 6 m to 19.6 μg/L at 18 m (Figure 1.10). In both lobes of Lake

Bonney, the hypolimnia were greater than two orders of magnitude more concentrated in Rb than their respective epilimnia: 2.80 μg/L to 294 μg/L in East Bonney and 1.50 μg/L to 156 μg/L in 16

1.5 0.20

1.2 0.16 g/L)

μ 0.9 0.12

0.6 0.08 Li:Cl (g:mg) Li:Cl Lithium ( Lithium 0.3 0.04

0.0 0.00 1.0 0.20

0.8 0.16 g/L) μ 0.6 0.12

0.4 0.08 Rb:Cl (g:mg) Rb:Cl

Rubidium ( Rubidium 0.2 0.04

0.0 0.00 32 4

24 3

g/L)

16 2

8 1 (g:mg) Sr:Cl Strontium ( Strontium

0 0

3.0 Lake Bull 0.6 2.5 0.5 Boulder Pavement Boulder 0.4

2.0 g/L)

μ 0.3 1.5 0.2

1.0 Ba:Cl (g:mg) Ba:Cl Barium ( Barium 0.1 0.5 6E-16 0.0 -0.1 0 5 10 15 20 25 30 35

Distance Downstream (km)

Figure 1.6 Longitudinal profile of the Onyx River. Closed symbols represent concentration in μg/L. Open symbols represent the minor ion concentration to chloride concentration ratio in g:mg.

17 0.4 0.40

0.3 0.30

g/L) μ

0.2 0.20 Li:Cl

Lithium ( Lithium 0.1 0.10

0.0 0.00 1.0 0.5

0.8 0.4 g/L) μ 0.6 0.3

0.4 0.2 Rb:Cl

Rubidium ( Rubidium 0.2 0.1

0.0 0 24 12 20 10

g/L) 16 8 μ

12 6 Sr:Cl 8 4

Strontium ( Strontium 4 2 0 0 4 4

3 3 g/L)

μ 2 2 Ba:Cl

Barium ( Barium 1

0 0 0 200 400 600 800 1000 Distance Downstream (m)

Figure 1.7 Longitudinal profile of Canada Stream. Closed symbols represent concentration in μg/L. Open symbols represent the minor ion concentration to chloride concentration ratio in g:mg.

18 West Bonney. The bottom water of Lake Vanda was nearly 38 times more concentrated in Rb than the surface water (94.3 μg/L and 2.50 μg/L, respectively). Don Juan Pond water also had ahigh concentration of Rb (102 μg/L), but its concentration was lower than the bottom waters of both lobes of Lake Bonney. Don Juan Pond had the lowest Rb:Cl value of the lakes sampled in this study (Figure 1.11). Lake Hoare and West Lake Bonney had fairly constant Rb:Cl values with depth, and Lake Fryxell and Lake Vanda had Rb:Cl values that decreased with depth. East Lake

Bonney had its lowest Rb:Cl values at the chemocline.

Strontium concentrations in the MCM lakes range from 1.5 x 102 μg/L in the surface water of Lake Hoare to 1.04 x 106 μg/L in Don Juan Pond (Appendix D). Lakes Hoare, Fryxell,

Bonney, and Vanda are stratified with respect to Sr concentrations (Figure 1.12). Although my data for Lake Bonney were limited to three data points in each lobe, the stratification patterns for the Taylor Valley lakes were in agreement with previously published work (JONES and FAURE,

1978; LYONS et al., 2002) Lake Hoare had the lowest concentration of Sr of the MCM lake surface waters (150 μg/L) followed by Lake Vanda (308 μg/L), West Lake Bonney (399 μg/L), East Lake

Bonney (624 μg/L), and Lake Fryxell (668 μg/L). Of the lake bottom waters, Lake Vanda had the highest concentration of Sr (6.93 x 104 μg/L) followed by West Lake Bonney (5.04 x 104 μg/L), East

Lake Bonney (4.51 x 104 μg/L), Lake Fryxell (7.74 x 103 μg/L), and Lake Hoare (928 μg/L). Lakes

Hoare and West Bonney had nearly constant Sr:Cl values with depth, and the Sr:Cl profile for

Lake Vanda was nearly constant (Figure 1.13). The Sr:Cl profiles for Lake Fryxell and East Lake

Bonney decreased with depth. Compared to the Sr:Cl profiles, Sr:Ca and Sr:Mg profiles showed little variation between lakes and throughout the water columns of the lakes (Figure 1.14, Figure

1.15). Sr:Ca ratios were fairly constant with depth in Lakes Vanda, Fryxell, and Hoare. Don Juan

Pond and the surface waters of Lakes Bonney, Hoare, and Fryxell were similar, but unlike Lakes

Vanda, Hoare and Fryxell, the Sr:Ca ratios in both lobes of Lake Bonney increased with depth. 19 The hypolimnia of Lake Bonney were the only samples with Sr:Ca values equal to or greater than seawater values. The Sr:Mg ratios were highest in Don Juan Pond, and Lake Hoare had higher

Sr:Mg ratios than Lakes Vanda, Bonney, and Fryxell. The Sr:Mg values in the water columns of

Lake Hoare, Lake Vanda and West Lake Bonney only decreased slightly with depth compared to

Lake Fryxell and East Lake Bonney. The deepest layers of West Lake Bonney and Lake Fryxell were similar to seawater values, and East Lake Bonney was the only lake with Sr:Mg ratios below the seawater ratio.

Barium concentrations in the lakes ranged from 2 μg/L in the surface water of Lake

Hoare to 1,400 μg/L in Don Juan Pond (Appendix D). The surface water of all the ice-covered lakes, Hoare, Fryxell, Bonney, and Vanda had Ba concentrations below 10 μg/L. Of these lakes,

Lake Vanda had the highest Ba concentration of 660 μg/L at 70 m, i.e. it was 82.5 times more concentrated in Ba than the surface water. The Ba concentrations of the hypolimnia in the other lakes were much higher than their epilimnia (Figure 1.16). The east (140 μg/L) and west (200

μg/L) lobes of Lake Bonney were 15.6 and 50 times more concentrated than their surface water concentrations. At 30 m, Lake Hoare water had a Ba concentration of 13 μg/L making it 7.5 times more concentrated than the surface water (2 μg/L). Lake Fryxell showed the smallest change in concentration with depth; the Ba concentration increased 6.6 fold from 5 μg/L at 6 m to 33 μg/L at

18 m. Lake Hoare had the highest Ba:Cl values as well as the highest K:Cl and Mg:Cl values

(Figure 1.17) (LYONS et al., 1998a). The hypolimnion of East Lake Bonney had the lowest value.

The Ba:Cl profile in Lake Vanda and both lobes of Lake Bonney decreased with depth. The profiles in Lake Hoare and Lake Fryxell showed that the Ba:Cl values remained fairly constant with depth.

20 1.5.4 Minor elements in lake sediments and suspended matter

Lithium concentrations in all the leached fractions of the lake sediments and suspended matter were very low and undetectable in many samples. All samples had at least trace amounts of Li in the sulfate phase, and no Li in the carbonate phase (Table 1.4). Only three samples from

West Lake Bonney (39 m), East Lake Bonney (38.4 m), and Lake Vanda had Li in the adsorbed phase (Figure 1.18). Total concentrations were highest in the surface sediments collected from the bottom of Lake Fryxell (3.30 µg/g) and from the suspended matter at 39 m in West Lake Bonney

(3.00 µg/g). In both of these samples, Li was dominantly in the sulfate phase. In East Lake

Bonney (38.4 m) and Lake Vanda, total concentrations were much lower (0.26 µg/g and 0.18 µg/g, respectively), and the majority of the Li was in the adsorbed fraction.

Rubidium concentrations in the lake sediments and suspended matter were also low

(Table 1.4). With the exception of West Lake Bonney at 10 m and East Lake Bonney at 37.5 m, both of which were not analyzed for the adsorbed or carbonate fraction, all samples had Rb exclusively in the sulfate phase (Figure 1.19). The concentrations ranged from 0.92 µg/g (Lake

Vanda) to 3.00 (Lake Fryxell at 18 m).

Li (μg/L) Li:Cl (g:g x 106) 1 10 100 1000 10000 100000 1000000 1 10 100 1000 0 0

10 10 Seawater

20 20

30 30

Depth (m) Depth (m) Depth

40 40

50 Hoare 50 Hoare Fryxell Fryxell E. Bonney E. Bonney 60 W. Bonney 60 W. Bonney Vanda Vanda Don Juan Pond Don Juan Pond 70 70

Figure 1.8 Lithium concentration by depth for the MCM lakes. Figure 1.9 Li:Cl values for the MCM lakes. Seawater from (PILSON, 1998).

22 Rb (μg/L) Rb:Cl (g:g x 106) 1 10 100 1000 0.1 1 10 100 0 0

10 10

20 20

30 30 Seawater

Depth (m) Depth (m) Depth 40 40

50 Hoare 50 Hoare Fryxell Fryxell E. Bonney E. Bonney 60 W. Bonney 60 W. Bonney Vanda Vanda Don Juan Pond Don Juan Pond 70 70

Figure 1.10 Rubidium concentration by depth for the MCM Figure 1.11 Rb:Cl values for the MCM lakes. Seawater from (PILSON, lakes. 1998).

23 Sr (μg/L) Sr:Cl (g:g x 106) 1 100 10000 1000000 100 1000 10000 0 0

10 10 Seawater

20 20

30 30

Depth (m) Depth (m) Depth 40 40

50 Hoare 50 Hoare Fryxell Fryxell E. Bonney E. Bonney 60 W. Bonney 60 W. Bonney Vanda Vanda Don Juan Pond Don Juan Pond 70 70

Figure 1.12 Strontium concentration by depth for the MCM Figure 1.13 Sr:Cl values for the MCM lakes. Seawater from (PILSON, lakes. 1998).

24 Sr:Ca (g:g x 103) Sr:Mg (g:g x 103) 1 10 100 1 10 100 1000 0 0

10 10 Seawater

20 20

30 30

Depth (m) Depth (m) Depth

40 40

Hoare Hoare

50 Fryxell 50 Fryxell E. Bonney E. Bonney W. Bonney W. Bonney 60 60

Vanda Vanda Seawater Don Juan Pond Don Juan Pond 70 70

Figure 1.14 Sr:Ca values for the MCM lakes. Seawater from Figure 1.15 Sr:Mg values for the MCM lakes. Seawater from (PILSON, (PILSON, 1998) 1998)

25 Sr (μg/L) Sr:Cl (g:g x 106) 1 100 10000 1000000 100 1000 10000 0 0

10 10 Seawater

20 20

30 30

Depth (m) Depth (m) Depth 40 40

50 Hoare 50 Hoare Fryxell Fryxell E. Bonney E. Bonney 60 W. Bonney 60 W. Bonney Vanda Vanda Don Juan Pond Don Juan Pond 70 70

Figure 1.16 Barium concentration by depth for the MCM lakes. Figure 1.17 Sr:Cl values for the MCM lakes. Seawater from (PILSON, 1998).

Sr concentrations were considerably higher than Li and Rb with total concentrations ranging from 2.40 µg/g in East Lake Bonney at 7 m to 1,600 µg/g in East Lake Bonney at 38.4 m

(Table 1.4). In Lake Fryxell (all depths), West Lake Bonney (38.7 m and 39 m), and East Lake

Bonney (38.4 m) the carbonate phase contained the most strontium (Figure 1.20). With the exception of East Lake Bonney at 7 m, strontium was substantially contained in the adsorbed phase. In West Lake Bonney (39 m), East Lake Bonney (7 m), Lake Hoare, and Lake Fryxell (all depths) strontium was present in the sulfate fraction.

Excluding the samples that were not analyzed for all phases, total Ba concentrations in the sediments and suspended matter ranged from 4.95 µg/g in East Lake Bonney (38.4 m) to 83.3

µg/g (Table 1.4). In East Lake Bonney, Ba was predominately in the adsorbed phase (Figure 1.21).

In West Lake Bonney, Lake Hoare, and Lake Fryxell, Ba was mostly contained in the operationally defined sulfate minerals. With the exception of East Lake Bonney, all samples that were analyzed for the carbonate fraction had substantial amounts of barium in the carbonate phase.

Stream sediments were previously collected from Taylor Valley and analyzed by XRF

(Table 1.5). The concentrations of all of these elements were much higher in the stream sediments than in the lake sediments and suspended matter. Unlike the lake sediments and suspended matter which had Sr in highest concentration, the stream sediments had Ba in highest concentration followed by Sr, Rb, and Li.

Li (µg/g) Rb (µg/g) Sr (µg/g) Ba (µg/g)

Sulfate Sulfate Sulfate Sulfate

Adsorbed Adsorbed Adsorbed Adsorbed

Carbonate Carbonate Carbonate Carbonate

WLB 10 DL n.d. n.d. DL n.d. n.d. 66.6 n.d. n.d. 11.0 n.d. n.d. WLB 39 0.15 2.9 DL DL 1.20 DL 192 144 241 8.21 14.0 15.4 WLB 38.7 DL 0.02 DL DL 1.50 DL 261 DL 290 2.55 12.6 11.8 28 ELB 7 DL 0.02 DL DL 0.60 DL DL 2.40 DL 15.5 DL DL

ELB 38.4 0.21 0.05 DL DL 1.20 DL 757 DL 847 3.68 1.28 DL ELB 37.5 DL n.d. n.d. DL n.d. n.d. 512 n.d. n.d. 0.08 n.d. n.d. HOR Surface DL 0.02 DL DL 1.80 DL 65.7 84.9 39.5 2.33 15.1 14.5 FRX 18 DL 0.03 DL DL 3.00 DL 49.7 87.2 146 14.1 42.8 26.4 FRX 19.4 DL 0.03 DL DL 1.80 DL 42.9 55.5 112 2.40 39.9 26.5 FRX Bottom DL 3.30 DL DL 2.40 DL 42.6 51.2 160 2.48 46.4 27.2 VAN 0.12 0.06 n.d. DL 0.92 n.d. 36.8 DL n.d. DL 4.15 n.d.

Table 1.4 Lithium, rubidium, strontium, and barium concentrations in the adsorbed, sulfate, and carbonate phases in MCM lakes.

Li (mg/g) Rb (mg/g) Sr (mg/g) Ba (mg/g) West Lake Bonney Basin 14 109 312 684 East Lake Bonney Basin 14 86 404 543 Lake Hoare Basin 16 70 582 719 Lake Fryxell Basin 15 78 504 690

Table 1.5 Average lithium, rubidium, and strontium concentrations of stream sediments analyzed by XRF.

4.0

3.0

2.0 Li (ug/g) Li

1.0

** 0.0 * *

Figure 1.18 Lithium concentration contained in the adsorbed (blue), sulfate (red), and carbonate (green) phases of MCM lake sediments. *Only adsorbed fraction analyzed. **Only adsorbed and sulfate fractions analyzed.

3.0

2.0 Rb (ug/g) Rb 1.0 **

0.0 * *

30 2000

1500

1000 Sr (ug/g) Sr

500 *

* 0 **

100

Ba (ug/g) Ba 40

20 * ** 0 *

1.6 Discussion

1.6.1 Minor alkali metals and alkaline earths in MCM streams

1.6.1.1 Solute acquisition in MCM streams

The source of solutes in the McMurdo Dry Valley streams is a combination of the chemistry of the snowfall in the accumulation areas of glaciers, chemical weathering/salt dissolution in supraglacial streams, and chemical weathering/salt dissolution in proglacial streams (LYONS et al., 1998a). Stream water in the MCM consists exclusively of glacial melt, thus the geochemical processes producing solutes on the glaciers is a fundamental component of stream geochemistry. The topography of the Dry Valleys, dominant east-west winds, abundant exposed soils, and proximity to features such as the East Antarctic Ice Sheet and McMurdo Sound make for complex patterns of glaciochemistry, however some generalizations can be made. For example, the chemistry of glaciers nearest the coast have relatively abundant amounts of sodium, chloride, sulfate, and calcium derived from marine aerosols and dust (WITHEROW et al., 2006).

Strong winds redistribute sediments and salts from the valley floor onto glacial surfaces and into supraglacial streams channels (FORTNER et al., 2005). As the temperatures rise during the summer months, glacial melt water solubilizes these salts and chemical weathering of silicates transported from the valley floor onto the glaciers contributes to the solute load in MCM streams

(FORTNER et al., 2005; LYONS et al., 1998a).

Despite the cold, arid conditions in the Dry Valleys, solute acquisition through chemical weathering occurs on the valley floor at considerable rates (GOOSEFF et al., 2002; LYONS et al.,

1997; MAURICE et al., 2002). There are several processes in cold deserts that are favorable to weathering. Frost wedging, salt weathering, freeze-thaw, and wind abrasion may expose fresh surfaces to be chemically weathered by water (UGOLINI, 1986). Although water in the McMurdo

Dry Valleys is spatially and temporally limited, there is a high degree of weathering where water 32 is present, and during periods of high discharge, chemical weathering processes have been shown to increase (GREEN et al., 2005; NEZAT et al., 2001). Particularly in the longer streams, the dissolution of evaporite salts and high-stand lacustrine carbonates and silicate weathering are significant processes in solute acquisition (GREEN et al., 2005; LYONS et al., 1998a). Chemical weathering in the hyporheic zone is also a substantial source of solutes to the MCM streams

(GOOSEFF et al., 2002; MAURICE et al., 2002).

1.6.1.2 Minor alkali behavior

The dominant form of transport of minor elements depends on the mobility of an element in the weathering and transport process and the amount of solids transported in a stream (GAILLARDET et al., 2003). Globally, the minor alkali metals and alkaline earths in streams are strongly correlated with either Na, Ca, or total dissolved solids (TDS) (EDMOND et al., 1996;

EDMOND et al., 1995). In the MCM streams, the minor alkali elements, Li, Rb, and Sr, correlate well with either Na, Ca, or TDS, but Ba shows no strong relationship with any of these.

Dissolution and leaching dominate early weathering stages whereas during advanced stages, exchange and adsorption onto clays are a significant process affecting alkali metals and alkaline earths in streams (NESBITT et al., 1980). In the MCM, chemical weathering in the stream channels can be either weathering or transport limited depending on the discharge and morphology of the stream (NEZAT et al., 2001). Given these conditions, it may be expected that as discharge varies greatly, especially temporarily, and the morphology of a stream channel changes

(e.g., the thawing and expansion of the hyporheic zone) the relative contributions from dissolution, leaching, cation exchange, and adsorption may vary temporarily and perhaps spatially as well. Recent on-going work shows that iron and manganese oxides and hydroxides are important components in MCM streams (FORTNER et al., submitted). Fe and Mn

33 oxides/hydroxides are very efficient absorbents of both cations and oxyanions depending on the pH (MOREL and HERRING, 1993). On high dielectric constant minerals such as manganese dioxide, the equilibrium constants for adsorbing alkaline earths are such that Mg2+ < Ca2+ < Sr2+ <

Ba2+, and on low dielectric constant minerals such as FeOH, the equilibrium constants of these elements have an opposite pattern (Mg2+ > Ca2+ > Sr2+ > Ba2+) (SVERJENSKY, 2006). It is evident that the composition of the suspended and bedload leaching experiments that absorption is fundamental in removing alkaline earths from solution in these streams, however more research needs to be conducted to quantify the amount of alkaline earths being removed by specific mineral phases.

Lithium is highly soluble forming stable complexes with water in tetrahedral geometry (HUH et al., 1998). This high solubility is illustrated by the strong relationship between Li and Cl, another highly soluble ion (Figure 1.2). Comparison of the average concentrations of Li, Rb, Sr, and Ba in stream sediments with the dissolved load in the streams reveals that the partitioning of

Li into solution is similar to Sr, but at least an order of magnitude higher than Rb and Ba (Table

1.5, Appendix C). For an alkali element, Li has a small ionic radius (0.78 Å) so it behaves much like Mg (0.72 Å), and like Mg, Li can undergo cation exchange reactions with clays (HUH et al.,

1998). For example, in pure experiments, Li is selectively adsorbed onto clay surfaces and fixed into non-exchangeable positions (ANDERSON et al., 1989). Lithium in major global rivers are positively correlated with magnesium, and this holds true for MCM streams (Figure 1.22) (HUH et al., 1998). Stream tracer experiments further corroborate the assertion that Li is being adsorbed onto clays. It has been shown that Li is preferentially taken up onto stream sediments while Na and K are released (GOOSEFF et al., 2004).

34 During rock-water interactions, Rb often completely partitions into the aqueous phase

(SEYFRIED et al., 1984). Because the larger Rb ion is preferentially retained over K on clay minerals, cation exchange is likely a dominant control on Rb concentration in streams (NESBITT et al., 1980). It is clear from the relationship between Rb and Cl that Rb is not conservative in these streams, and it is likely being removed from solution (Figure 1.3). The correlation between K and

Rb in these stream waters is poor, indicating that the removal of these cations by clay adsorption is complex. The Rb/K relationship with Cl in streams sheds light on the uptake mechanisms in the streams. At low chloride concentrations (i.e. less than 5 mg/L), Rb/K values are highly variable (Figure 1.23). At Cl concentrations greater than 5 mg/L, nearly all Rb/K values are less than 1 mg/g (Figure 1.24). This pattern may suggest that in stream water above a threshold of 5 mg/L Cl, Rb is removed from stream water while K remains in solution.

1 Li (ug/L) Li 0.1

y = 0.66x + 0.22 R² = 0.81

0.01 0.1 1 10 Mg (mg/L)

Figure 1.22 Magnesium and lithium concentration in MCM streams.

y = 0.73x0.34 4 R² = 0.27

g/L) μ

Rb ( Rb 2

0 0 1 2 3 4 5 K (mg/L)

Figure 1.23 K to Rb relationship in stream water.

Rb/K (g/mg) Rb/K 2

0 0 5 10 15 20 25 30 35 40 Cl (mg/L)

Figure 1.24 Cl versus K/Rb relationship in stream water.

The concentration of strontium, particularly using the isotope ratio, Sr87/Sr86 is a classis tracer of rock weatheirng (FAURE and MENSING, 2005; GAILLARDET et al., 2003). Strontium is rapidly and strongly solubilized during the weathering of fresh continental rocks, and during chemical weathering, Sr and Ca are removed from silicates and carbonates at the same rate

(NESBITT et al., 1980). In MCM streams, this is evident in the relationship between strontium and calcium (Figure 1.25). Previous work on the strontium isotopes in the Onyx River has shown that the 87Sr/86Sr values are similar to those of soil in Wright Valley suggesting that interactions between soils/sediments and streams play a significant role in mobilizing Sr (JONES and FAURE,

1967). Research conducted in Taylor Valley has also concluded that chemical weathering is occurring in the dry valley streams and is solubilizing strontium (LYONS et al., 2002). This is further evidenced by relatively high ratio of Sr in stream water and sediments compared to Rb ad

Ba (Table 1.5, Appendix C). In streams where cation exchange with clay minerals is occurring, deviations from an ideal Sr to Ca relationship is likely due to the preferential removal of Sr relative to Ca from solution (NESBITT et al., 1980). Geochemical modeling of streams indicates that clay minerals such as smectites, illites, and vermiculites may be produced in the stream channel, and empirical evidence has shown that smectite is present in wetted soils (CAMPBELL and CLARIDGE, 1982; LYONS et al., 1998a). Because the relationship between Sr and Ca concentrations in the streams is so strong (r2 = 0.92), preferential uptake of Sr from clays is likely limited.

K-feldspar is considered to be the predominant source of Ba in granitic rocks, however weathering profiles reveal that Ba may either be released into solution or retained on clay minerals (NESBITT et al., 1980; PUCHELT, 1972). In basaltic rocks, it is the least mobile of the alkaline earth elements possibly due to weathering-resistant minerals, or sequestration on clay minerals or Fe-oxyhydroxides (DAS and KRISHNASWAMIN, 2006). The poor relationship between 37 K and Ba in MCM stream waters may indicate that Ba is not being consistently removed from silicate rocks due to simplified chemical weathering processes (Figure 1.26). In addition to chemical weathering of silicates, the dissolution of carbonates and salts may be contributing Ba to the dissolved load of the streams. It is clear from the poor relationship between Cl and Ba, that

Ba is not behaving conservatively in the streams, that is, it is being removed from solution

(Figure 1.5). Although most alkali metals and alkaline earths are predominantly partitioned into the dissolved phase in aqueous systems, considerable amounts of barium have been observed to be in the suspended phase in some rivers (TANIZAKI et al., 1992).

1000

100

g/L)

μ Sr ( Sr 10

y = 5.96x + 2.07 R² = 0.92 1 0.1 1 10 100 Ca (mg/L)

Figure 1.25 Sr to Ca relationship in MCM stream water.

38 16

10 g/L)

μ 8

Ba ( Ba 6

0 0 1 2 3 4 5 6 K (mg/L)

Figure 1.26 Potassium to barium concentrations in MCM stream water.

1.6.1.3 Stream transects

The length of a stream is an important variable in solute content of streams as longer streams in the MCM typically have higher total dissolved solids and greater Ca:Na and HCO3:Cl values (GREEN et al., 1988). It has also been shown that the Onyx River, for example, goes from a

“precipitation dominated” stream at its headwaters to “rock dominated” downstream (GREEN and CANFIELD, 1984). For example, the headwaters of the Onyx River have Ca:Na and Cl:HCO3 ratios similar to direct melt from the Wright Glacier, and the concentrations of major ions and Si increase with increasing distance downstream indicating that chemical weathering is occurring in the streams (GREEN et al., 2005; MAURICE et al., 2002)

Longitudinal increases in concentration in the MCM streams for Li, Rb, and Sr indicate that weathering and dissolution of salts are contributing to the increase in downstream solutes

(Figure 1.6, Figure 1.7) (GOOSEFF et al., 2002; NEZAT et al., 2001). In the Onyx River, lithium and strontium concentrations rapidly increase after Bull Lake, indicating that this section of the

39 stream may be a source of lithium and strontium (Figure 1.6). Other processes besides local weathering such as rapid particulate scavenging and recycling of pore waters contribute to longitudinal variation in trace metal concentrations in the Onyx River and these processes may also increases other solute chemistry (GREEN et al., 2005).

Previous work identified a section of the Onyx River lined with a “boulder pavement” as a nutrient sink as nitrate and phosphate concentrations decrease between 29.1 and 31.5 km downstream (Figure 1.6) (GREEN et al., 2005; HOWARD-WILLIAMS et al., 1997). This boulder pavement is a 1.5 km stretch of the stream channel with abundant flat boulders, algae, and cyanobacteria (HOWARD-WILLIAMS et al., 1997). Downstream from the boulder pavement is Bull

Lake, a minor NO3 sink (GREEN et al., 2005). Barium has been use as a proxy for marine productivity and it may possibly be taken up by cyanobacteria or algae mats in these locations in the Onyx River with abundant biomass (BERNSTEIN and BYRNE, 2004). However, although nutrients are taken up by the biota in the boulder pavement and removed from the stream in Bull

Lake, there appears to be no relationship between Ba and NO3 in the Onyx River and Canada

Stream (LYONS, 2009). My results show that while the Ba concentration decreases slightly, there is much variability in Ba concentration in the length of the Onyx River making it difficult to determine if there is a sink for barium (Figure 1.6). Lithium and Sr concentrations rapidly increase after the boulder pavement indicating that this section of the stream may be a source of

Li and Sr, but it is clearly not a significant source or sink of Rb since the concentration remains nearly the same.

1.6.2 Lake chemistry

The age of the lakes, location in the MCM, and sensitivity to climate change have influenced their geochemical evolution (LYONS et al., 2000; LYONS et al., 2005; LYONS et al., 1998a).

40 If inflow waters have a (HCO3 + CO3):Ca (M:M) value less than 2, calcium-enriched, carbonate- depleted lakes evolve from streams (Figure 1.27) (HARDIE and EUGSTER, 1970). Likewise if stream waters have (HCO3 + CO3):Mg (M:M) value less than 2, magnesium-enriched lakes evolve. MCM streams show considerable variation in alkalinity to calcium ratios (Figure 1.28). This variability is due to the source of these solutes to the streams (e.g. the dissolution of carbonates, silicate weathering, or the same origins of Blood Falls) (GREEN et al., 1988; LYONS et al., 1998a; NEUMANN,

1999). Lakes Fryxell and Lake Hoare should evolve to HCO3-rich waters based on the first step of geochemical evolution: the precipitation of calcite (GREEN et al., 1988; LYONS et al., 1998a). At the next geochemical divide after all Ca has been depleted, Lakes Fryxell and Hoare diverge in their evolution and the surface waters of Lake Fryxell would be dominated by Na-HCO3-CO3, and

Lake Hoare would be a Na-Mg-Cl-SO4-rich brine (LYONS et al., 1998a). The surface waters of

Lake Bonney would also become a Na-Mg-Cl-SO4-rich brine like Lake Hoare, but from a different path (LYONS et al., 1998a). Evaporation of Onyx River water would have resulted in a Na-Mg-Cl-

HCO3 brine rather than the Ca-Na-Mg-Cl-rich waters in the hypolimnion of Lake Vanda today

(GREEN and CANFIELD, 1984). Although these lakes should evolve towards these brines, this and previous data show that rather than the brined being depleted in Ca and Mg, they are enriched relative to HCO3 + CO3 (Figure 1.27). Only Lake Fryxell is depleted in Ca. These depletions are relative as when the Ca and Mg concentrations are compared to Cl concentrations, it is evident that Ca and Mg are being removed in the lakes as well (Figure 1.28).

1.6.2.1 Lake Hoare

The primary source of solutes to the Lake Hoare is the dissolution of marine aerosols, aeolian salts, and chemical weathering of dust on Canada Glacier in addition to inputs from

Andersen Creek and overflow from Lake Chad (LYONS et al., 2005; TEGT, 2002). In addition to

10000

1000

100 HOR FRX 10 ELB WLB 1 VAN

Streams Ca (mM),Ca Mg (mM) 0.1 DJP

0.01

0.001 0.01 0.1 1 10 100 HCO3+CO3 (mM) 1

0.9

0.8

0.7

0.6

0.5 2:1

0.4

Ca (mM),Mg Ca (mM) 0.3

0.2

0.1

0 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1

HCO3+CO3 (mM)

Figure 1.27 Bicarbonate + carbonate concentration versus calcium and magnesium concentration for the MCM lakes and streams. In addition to the log-log plot, stream data are also represented on a linear plot to show more detail. Open symbols represent calcium concentrations, and solid symbols represent magnesium concentrations. Dashed line is a 1:1 ratio. HOR = Lake Hoare; FRX = Lake Fryxell; ELB = East Lake Bonney; WLB = West Lake Bonney; VAN = Lake Vanda; DJP = Don Juan Pond. Although the bicarbonate + carbonate concentration Don Juan Pond data are below the detection limit, it is plotted along the y-axis such that its calcium and magnesium concentrations can be compared to the streams and other lakes.

10000

1000

100 HOR FRX 10 ELB WLB 1 VAN Ca (mM),Mg Ca (mM) Streams 0.1 DJP

0.01

0.001 0.001 0.1 10 1000 100000 Cl (mM)

this work, earlier work on the Sr:Ca values of Lake Hoare and other MCM lakes shows that the ratios are well below the natural variability in the oceans (Lyons et al., 2002). This suggests that either the source of alkaline earth elements in the valleys is enriched in Ca or that

Ca is preferentially solubilized during chemical weathering. Analysis of solutes in glacial snow shows that wind-blown dust is a dominant source of Ca and Cl to the glacier surfaces and that the dissolution of dust is also an important source of other solutes, particularly those associated with crustal material(Witherow et al., 2006). The Li:Mg ratio in the water column of Lake Hoare shows that Li and Mg have a common source, and that this source is unlike seawater (Figure

1.29). The Li: Cl, Rb:Cl, Sr:Cl, and Ba:Cl values for Lake Hoare are all higher than the seawater ratios (Figure 1.9, Figure 1.11, Figure 1.13, Figure 1.17). This further illustrates that the present- day marine-derived aerosol or glacial melt is not the sole source of Li, Rb, Sr, and Ba.

1.6.2.2 Lake Fryxell

The geochemistry of Lake Fryxell is dominated by the diffusion of some solutes (e.g. Cl) from a brine beneath the sediment water interface and the infilling of the lake by recent glacial melt containing chemical weathering and marine salt components (LYONS et al., 2005). The Li:Cl and Ba:Cl values for Lake Fryxell are higher than those in seawater (Figure 1.9, Figure 1.17). At the end of the last draw-down event (approx. 1 ka), it has been proposed that Lake Fryxell was saturated with respect to NaCl (LYONS et al., 1997). Because of the higher solubility of Li, Cl was relatively depleted resulting in an increase of the Li:Cl ratio in the bottom of Lake Fryxell (LYONS and WELCH, 1997). Lithium isotopic evidence discussed in Chapter 2 further argues the case for this marine source.

44 Above 12 m, Lake Fryxell has higher Rb:Cl values than seawater, and below 12 m, the

Rb:Cl values in the lake are very similar to seawater (Figure 1.11). This pattern suggests that Rb is strongly influenced by diffusion of marine salts from the sediment-water interface and the Rb in the surface water is from a different source: chemical weathering and salt and dust dissolution as with Lake Hoare.

In the bottom of Lake Fryxell, the Sr:Cl values are very similar to the ratio in seawater corroborating the assertion made by others that the bottom of Lake Fryxell is heavily influenced by seawater (Figure 1.13) (LYONS et al., 2005). The Sr:Cl values progressively increase towards the surface of the lake which signifies that there is an additional source of Sr to the lake. The inflow waters to Lake Fryxell have 87Sr/86Sr ratios that are all slightly more radiogenic than the surface water of the lake demonstrating that the present sources of strontium do not reflect sources of strontium to the lake in the past (LYONS et al., 2002). Strontium isotopes suggest that weathering of volcanic rocks has contributed a significant amount of Sr to Lake Fryxell (LYONS et al., 2002)

1.6.2.3 Lake Bonney

The two lobes of Lake Bonney are remnants of ancient marine waters that have been modified by chemical weathering inputs (LYONS et al., 2005). The unusual chemistries of the lobes of Lake Bonney are in response to climatic changes in the MCM over at least the past

300,000 years, as the lake shows evidence that it has lost solutes during cryogenic concentration and mineral precipitation (LYONS et al., 2005). In both lobes of Lake Bonney, the Li:Cl values are higher than that of seawater (Figure 1.9). Although lithium may have several sources in the Lake

Bonney Basin, the ratio of Li:Cl is constant with depth, indicating that through time, the sources of lithium to the lake have remained relatively unchanged. However, earlier work on the Li

45 concentrations in Lake Vanda examined the ratio of Li to Na to conclude that climatological changes in the dry valleys has changed the source of lithium through time (LYONS and WELCH,

1997). Considering the advanced stage of evolution in the bottom brine of East Lake Bonney, Na is potentially being removed as Na-sulfate or Na-chloride minerals. Therefore, Li:Cl values may be a more useful indicator in determining the source of Li.

Li:Mg (g:g x 103) 0.1 1 10 0

10 Seawater

Depth (m) Depth 40

Hoare 50 Fryxell E. Bonney W. Bonney 60 Vanda Don Juan Pond 70

Figure 1.29 Li:Mg ratios in the MCM lakes. Note that Don Juan Pond is not included on this plot to better examine the variation in the other lake profiles. Seawater data (Pilson, 1998)

The west lobe of Lake Bonney has Sr:Cl values above seawater, and at and below the chemocline, the east lobe has ratios below seawater (Figure 1.13). Strontium in the surface of

West Lake Bonney is dominated by Blood Falls, and the streams flowing from Taylor Glacier,

Lyons Creek and Santa Fe Creek, have similar strontium isotope ratios to Blood Falls as well

(LYONS et al., 2002). The Sr isotopes in the bottom waters of Lake Bonney are unlike Blood Falls indicating that the strontium in these brines is not totally derived from Blood Falls or ancient seawater (LYONS et al., 2002).

The Ba:Cl values in both lobes of Lake Bonney are higher than seawater, but the hypolimnion of the east lobe is only slightly higher than seawater suggesting that the Ba in the bottom of East Bonney may be of marine origin (Figure 1.17). The Rb:Cl values in Lake Bonney are lower than seawater indicating that seawater may not be the dominant source of Rb to the lake, or that Rb has been removed by some mechanism (Figure 1.11). Perhaps a better explanation may be that Rb is being removed by adsorption onto sediments as shown in the sediment leaching experiments (Figure 1.19).

1.6.2.4 Lake Vanda

In Lake Vanda, the Li:Cl, Sr:Cl, and Ba:Cl values are higher than seawater indicating that seawater is not the only source of Li, Sr, and Rb to the lake currently and historically (Figure 1.9,

Figure 1.13, Figure 1.17). Early work on strontium isotope ratios in the Onyx River/Lake Vanda system indicates that a direct marine source of strontium is unlikely, but rather, rock weathering of igneous and metamorphic rocks is the likely source (JONES and FAURE, 1978). It has also been shown that the strontium in Lake Vanda could have come solely from inputs from the Onyx

River (JONES and FAURE, 1967). The surface water of Lake Vanda has a slightly higher Rb:Cl value than seawater, but at depth, this ratio decreases becoming lower than seawater (Figure

47 1.11). Such a pattern may indicate that Rb is not conservative in Lake Vanda, i.e. it is being removed from solution or has been removed in the past during the brine-generating low stands

(WILSON, 1964).

1.6.2.5 Don Juan Pond

The Li:Cl, Sr:Cl, ad Ba:Cl values in Don Juan Pond are much higher than in seawater corroborating other evidence that the pond’s water is not of marine origin (Figure 1.9,Figure 1.13,

Figure 1.17). Rb:Cl values are over an order of magnitude lower in Don Juan Pond water compared to seawater perhaps further corroborating a solute source other than oceanic (Figure

1.11). Don Juan Pond is underlain by silty sands atop the Ferrar dolerite (HARRIS and

CARTWRIGHT, 1981). The Ferrar Dolerite is a Rb enriched mafic sill that extends along the

Transantarctic Mountains (FLEMING et al., 1995). Both K and Rb are readily mobilized during chemical weathering, however the Rb:K ratios of Don Juan Pond (0.7 mg:g) are much lower than the average composition of sandstone (5.6 mg:g) and the Ferrar Dolerite (10.0 mg:g) (FLEMING et al., 1995; SEYFRIED et al., 1984). As seen in the stream chemistry, at chloride concentrations greater than 5 mg/l, Rb:K ratios decrease suggesting that Rb is being lost and possibly adsorbed onto clays (Figure 1.24). The low Rb:K ratio in Don Juan Pond may be the result of very high rock:water ratios of the incoming groundwater being retained in the aquifer materials. The origin of Don Juan Pond solutes will be further discussed in Chapter 2.

1.6.3 Geochemical sinks

1.6.3.1 Evaporites

During brine evolution, the primary process controlling the chemical composition of a closed basin lake is the precipitation of a series of minerals through evaporation (HARDIE and

48 EUGSTER, 1970). Much work has been done on the major ions, and it is clear that the precipitation of calcite and gypsum play a major role along with sepiolite, Mg-silicates, and magnesite in the evolution of closed-basin systems (EUGSTER and HARDIE, 1978; HARDIE and EUGSTER, 1970). The relative proportion of calcium and magnesium and their concentrations relative to HCO3 and SO4 play a critical role in the precipitation of these minerals, yet the behavior of the minor alkaline earth elements, strontium and barium, in closed basin lakes is largely unknown. These elements are capable of producing sulfate and carbonate minerals which may ultimately affect the composition of a brine. This idea of minor elements being removed by simple salt formation is further discussed in Chapter 2.

Sediment leaching experiments show that sulfates and carbonates are responsible for removing substantial quantities of Sr and Ba from the lakes (Figure 1.20, Figure 1.21). Because of strontium’s ability to substitute for Ca in carbonates, especially aragonite, it is not surprising to find that Sr is primarily found in association with the carbonate fraction of the sediments particularly in Lakes Bonney and Fryxell. Additionally, Sr may be precipitating as strontianite

(SrCO3) contributing the strontium content in carbonate sediments. Barite is mostly found in the sulfate fraction of the lake sediments, with the exception of East Lake Bonney. The specifics of sulfate precipitation in the lake is discussed in further detail in Chapter 2, but the likely form of the barite is in barium (BaSO4).

1.6.3.2 Clay absorption

Because of the extreme solubility of their simple salts, the alkali metals (Li and Rb) are not easily removed during salt precipitation. However, they may be adsorbed to negatively charged layers of authigenic clays (JAMES and PALMER, 2000). Lithium does not occur as a pure evaporate in concentrations greater than few ppb, rather it is found in association with

49 terregenous clay interactions (Boyko, 1966). It has also been shown that in the marine system, lithium uptake onto clays is an important transport mechanism delivering lithium to the ocean floor (STOFFYN-EGLI and MACKENZIE, 1984). Likewise, rubidium does not produce abundant salts, but instead may be removed by fixation onto the silicate layers of clays such as vermiculite

(AMPHLETT, 1964). Pore fluid data from marine sediment cores show that Li and Rb are transferred into marine sediments during early diagenesis, but it is not clear that marine sedimentation is a long-term sink particularly in the presence of high pore water concentration

NH4 (JAMES and PALMER, 2000). Due to the chemical stratification frequently observed in closed- basin lakes, an anoxic layer may form in the bottom strata of some of these lakes (e.g. like in Lake

Fryxell). It has been shown that under anoxic conditions, Li and Rb behave conservatively, and

Lake Fryxell shows no evidence of them being resolubilized (VIOLLIER et al., 1995). My leaching experiments show that little Li and no Rb is adsorbed to sediments suggesting that if adsorption is occurring, these alkali metals are being resolubilized and the sediments are not a permanent sink (Figure 1.18, Figure 1.19). On the other hand, the alkaline earth elements, Sr and Ba appear to be in significant quantities in the adsorbed fraction of the sediments indication that adsorption onto clays may an important removal mechanism.

1.6.3.3 Reverse weathering

Early work in the MCM has suggested that Lakes Fryxell and Hoare have K deficits that cannot be explained by the precipitation of evaporites, but may be the result of reverse weathering process(GREEN et al., 1988). These reactions initially put forth by (MACKENZIE and

GARRELS, 1966) are put simply as:

K-poor Al Silicates + HCO3- + SiO2 + K+ = K Al Silicates + CO2 + H2O

50 Such processes result in the removal of cations and alkalinity and the production of clay minerals and acidity. This reconstitution of silicates is thought to require anoxic sediments, silica, Fe- oxide, and organic matter (MACKENZIE et al., 1981). The hypolimnion of Lake Fryxell might be ideal for these reactions, and reverse weathering could provide an explanation for Li and Ba removal in the lake. There is little evidence to suggest that cations are being removed in the open ocean by this process, perhaps due to the low pH of the deep ocean (approximately 7.7)

(WOLLAST et al., 1968). Additionally, the hypersaline brine of the Orca Basin in the Gulf of

Mexico despite seemingly ideal conditions (SCHIJF, 2007). The closed basin lakes in the Ethiopian

Rift are Na-Cl-HCO3 type waters of moderate salinity, and work has shown that rapid Mg removal is being removed in association with clays, not through carbonate precipitation

(VONDAMM and EDMOND, 1984). It has been speculated that the high pH (> 9) in these lakes provide favorable conditions for reverse weathering. If a high pH is indeed a criterion for reverse weathering reactions, it is unlikely that the lakes of the MCM are producing silicates in their sediments. However, experiment have sucessfully shown that reverse weathering processes may be occurring in the muddy delta of the Amazon River where pHs are circum-neutral

(MICHALOPOULOS and ALLER, 1995). In order to thoroughly investigate the possibility of reverse weathering in the MCM, in-depth sedimentological experiments need to be conducted.

1.7 Conclusions

The major ions, Na, Ca, K, Mg, Cl, SO4, HCO3 + CO3 have been studied extensively in the

McMurdo Dry Valleys and in other closed basin lakes around the world. Despite their ability to accumulate in considerable concentrations, the minor elements Li, Rb, Sr, and Ba have been poorly studied. Here, I have presented my findings on the chemical composition of minor alkaki

51 metals and alkali earth elements in the streams and lakes of the McMurdo Dry Valleys. This work has resulted in the following conclusions:

 Lithium has a strong correlation with chloride in the streams indicating that

lithium is not being significantly removed from the streams.

 Rubidium and barium have complex behaviors in streams, perhaps being

removed through adsorption onto clay minerals.

 The marine salts are not the sole source of Li, Rb, Sr, and Ba to the lakes.

Chemical weathering likely contributes a significant amount of minor cations to

the streams and lakes.

 Small amounts of Li may be removed by sulfate precipitation in the bottom

water of Lake Fryxell and West Lake Bonney. Some Rb may be removed by

sulfate precipitation in the lakes. Sr and Ba appear to be removed in

considerable amounts in the lakes by adsorption onto clays, sulfate precipitation,

and carbonate precipitation.

1.8 Acknowledgements

I wish to thank W.B. Lyons, A. Carey, O. Costa, and D. Porinchu, at The Ohio State

University and W. Green at Miami University for their scientific and editorial assistance. K.

Welch and C. Gardner from the Byrd Polar Research Center aided in analyzing samples in

Antarctica. My thanks go to J. Priscu for the suspended matter samples and to P. Doran for the

Lake Frxyell sediment. J. Olesik, and A. Lutton from TERL at The Ohio State University provided assistance on the ICPMS. C. Westervelt provided her assistance conducting the sediment leaching experiments. J. Barker and researchers from the University of Colorado Stream Team aided in sample collection. This work was supported by NSF-ANT grant #0423595

Chapter 2 Lithium Isotopic Composition of the McMurdo Dry Valleys Aquatic Systems

2.1 Abstract

The McMurdo Dry Valleys of Antarctica are the largest ice-free area on the continent and include several terminal lake basins. These include Lakes Hoare, Fryxell, Bonney, and Vanda.

The hydrology of the lakes is controlled by the inputs of glacial meltwater and loss by sublimation of ice covers. Despite the extremely dilute inflow waters, the lakes range in salinity from hypersaline to fresh due in large part to their landscape position and climatic history.

Although the aqueous geochemistry of the region has been studied extensively over the past four decades, the ultimate source of solutes to the lakes is still in question. Atmospheric input/precipitation, dissolution of salts in sediments, chemical weathering, remnant seawater, and groundwater inputs have been suggested as the most likely sources of solutes to dry valley waters. I examined the lithium isotopic composition of glacial snow, stream water, and lake water to better understand the role of the above-mentioned processes in determining the overall chemistry of the lakes. My data suggest that the dominant sources of lithium to the lakes are the dissolution of aeolian dust of crustal origin on the glacial surface, marine aerosols, dissolution of salts and chemical weathering in the stream channels, and ancient seawater. The streams in the valleys show mixing of seawater (and/or marine aerosol) with an isotopically lighter glacial snow end-member. Data from Lake Hoare also reflect this trend suggesting inputs of lithium from these two sources. The monimolimnion of Lake Fryxell indicates a distinct marine end-member.

The lithium in Lake Bonney appears to be the result of mixing of a hypersaline brine from

53 beneath the Taylor Glacier with water from streamflow. The Lake Vanda system has experienced a high degree of chemical evolution, and the surface waters of the lake are a mixture of water from the Onyx River and diffusional input from the brine at the bottom of the lake. The lithium isotopic data suggest that the source of the lithium to Don Juan Pond, a hypersaline groundwater discharge playa, is from a deeper groundwater source of unknown origin that has been modified through rock-water interaction. This work provides a more conclusive link than previous findings between various salt sources and the role they play in solute accumulation in this extreme environment.

2.2 Introduction

Closed-basin, or terminal lakes occur on every continent and have been used to investigate modern and past geochemical processes (EUGSTER and HARDIE, 1978; JONES and

DEOCAMPO, 2004). These hydrologically closed lakes do not drain and are not flushed by the outflow of water, but rather lose water primarily through evaporation. Water loss from evaporation can cause these lakes to increase in salinity through time. This characteristic results in the precipitation of less soluble minerals and the concentration of residual waters rich in so- called, more soluble, conservative elements (EUGSTER and HARDIE, 1978; EUGSTER and JONES,

1979). Lithium has the highest ionic potential of all the alkali metals making it extremely soluble in aqueous solutions, and it is not easily accommodated in common mineral precipitates.

Because lithium is so soluble, it is readily transported in streams and behaves conservatively in lakes. Lithium isotopes are therefore useful in examining the sources of lithium, and by analogy, other soluble cations in modern hydrological and geochemical processes in these systems.

In addition to its high mobility in geologic settings, lithium has the potential to demonstrate significant isotopic fractionation due to the relatively large mass differences

54 between its two isotopes, 6Li and 7Li, making it a useful weathering tracer (HUH et al., 2001). It has been well established that because of its ability to form stronger chemical bonds, 6Li is preferentially retained in the solid phase and 7Li is readily mobilized (TOMASCAK, 2004).

Although there has been no definitive relationship identified between lithology and δ7Li, the δ7Li value of a stream generally reflects the isotopic signature of the local geology (HUH et al., 1998).

For example, streams draining black shales are +26‰, those draining shields are +6.6 to +19.4‰, and those draining mixed silicious terrains are +6 to +28.8‰ (HUH et al., 1998). The weathering character of a catchment influences the δ7Li as well. The lithium isotopes in the dissolved load of a river are heavier in reaction-limited, high-relief locales compared to locations where weathering reactions are driven to completion (HUH et al., 2001). In all streams, the δ7Li of the dissolved load is heavier than the suspended load by +5 to +30‰ (HUH et al., 2001; HUH et al., 1998).

Authigenic clays preferentially incorporate 6Li during their formation and thus cause this observed fractionation (HUH et al., 2001). These recent advancements in lithium isotope analysis present new opportunities for examining geochemical processes in hydrological systems.

The sources of solutes to the McMurdo Dry Valley (MCM), Antarctica lakes have been investigated for over four decades. Differentiation of these sources has been aided by numerous isotopic techniques: 87Sr/86Sr, 36Cl, δ37Cl, δ13C, 234U/238U (HENDERSON et al., 2006; JONES and FAURE,

1978; LYONS et al., 2002; LYONS et al., 1998b; LYONS et al., 1999b; LYONS et al., 1998c; NEUMANN et al., 1998). Angino and others (ANGINO et al., 1962) first suggested that there exist four potential sources of solutes to these lakes: chemical weathering of silicates in the stream channels, marine aerosol from glacial melt, dissolution of salts on the valley floor, and groundwater. Chemical weathering of the stream channel and hyporheic zone coupled with the dissolution of marine- derived salts are thought to be the dominant sources of lithium and other cations to the

McMurdo Dry Valley lakes (GOOSEFF et al., 2002; LYONS and WELCH, 1997; MAURICE et al., 2002). 55 The saline and hypersaline lakes are highly stratified, and, with perennial ice covers, little to no wind mixing occurs. The chemical profiles are dominated by diffusional processes (SPIEGEL and

PRISCU, 1998). Therefore, the chemical profiles of these lakes represent distinct environmental conditions and/or evolutionary histories that have influenced the sources of solutes to the lakes

(LYONS et al., 2005). Indeed, the lithium concentrations in Lake Bonney are among the highest in the world (LYONS and WELCH, 1997). The Li:Na profiles of the lakes vary with depth suggesting the potential for different sources of lithium to the lakes through time (LYONS and WELCH, 1997).

It has been proposed that the salts in Lakes Bonney and Fryxell originated from seawater and have since been modified by glacial melt and chemical weathering inputs (LYONS et al., 2005;

TAKAMATSU et al., 1998). Here, I use lithium isotope analysis to discuss the sources of solutes to the McMurdo Dry Valley lakes.

2.3 Site Description

The McMurdo Dry Valleys (76°30’-78°30’ S, 160°-164° E) are the largest ice free area in

Antarctica at approximately 4,800 km2. Taylor Valley and Wright Valley are situated in the

Trans-Antarctic Mountains approximately 10 km from the Ross Sea coast (Appendix E). The geology of the dry valleys has been described in detail and will not be repeated here (DORAN et al., 1994). Briefly, the geology consists of a Precambrian to Cambrian metamorphic and granitic basement rocks intertwined with felsic to intermediate igneous intrusions capped by the Beacon

Sandstone (mid-Paleozoic to mid-Mesozoic) in Wright valley and in some locations in Taylor

Valley (HASKELL et al., 1965). The McMurdo Volcanic Group rocks are located along the walls and on the floor of Taylor Valley, and may contain large inclusions of Jurassic Ferrar Dolerite

(WRIGHT and KYLE, 1990). The valley has been glaciated numerous times, and as such, glacial drift, tills, moraines, and lacustrine deposits cover the valley floor. The composition of these tills

56 consists of fragments of the surrounding local rock types: gneiss, diorite, schist, granite, diabase, sandstone, dolerite, and the more recent McMurdo volcanics ranging in age from 3.7 x 103 years to approximately 3.5 x 106 years (BOCKHEIM, 1997; HALL and DENTON, 2005). Soil development and composition in Taylor Valley is related to the landscape age such that the more developed soils and weathering features occur at older, higher elevations (BOCKHEIM, 2002). Efflorescent salts and encrustations occur at or below the soil surface (KEYS and WILLIAMS, 1981). The soils have a high salt content and are enriched in SO4, Cl, NO3, Na, K, Mg, and Ca (CLARIDGE and

CAMPBELL, 1977). Calcite, halite, gypsum, mirabilite, and other minerals occur in the soils of the dry valleys (KEYS and WILLIAMS, 1981). Much of the CaCO3 may be lacustrine in origin and was deposited when lake levels were much higher than today (HENDY et al., 1979). Clay minerals also exist in the soils, but smectite is thought to be the only Mg-rich authigenic clay mineral formed in the valleys (CAMPBELL and CLARIDGE, 1987).

Streams in the MCM are fed exclusively by glacial melt and flow for approximately 6-10 weeks during the austral summer (CONOVITZ et al., 1998). Less than 1 cm water equivalent of snow reaches the valley floor on any given year, and the snow that falls often accumulates in drifts in the stream channels, inhibiting flow (LYONS et al., 2000). In addition to snow dams, the low angle of the sun causes highly variable flow both daily and seasonally (CONOVITZ et al.,

1998). The length of the stream is an important variable on the acquisition of solutes. For instance, longer streams have higher total dissolved solids and evidence of more chemical weathering (GREEN et al., 1988). Chemical weathering rates within the hyporheic zone of the streams are quite high (GOOSEFF et al., 2002; MAURICE et al., 2002).

Most of the larger lakes in the McMurdo Dry Valleys are closed-basin, and lose water only through the sublimation of their permanent ice covers. The 3-5 m thick perennial ice results from the balance between the freezing of water at the bottom of the ice pack and sublimation at 57 the surface (CHINN, 1981; LYONS et al., 2000). During the summer months, a “moat” can form around the perimeter allowing glacial melt water to enter. Don Juan Pond is a hypersaline groundwater discharge playa with measured salinity of more than 40 wt% (HARRIS and

CARTWRIGHT, 1981). Due to its very high salinity, Don Juan Pond does not have an ice cover and therefore loses its water by evaporation. In the ice-covered lakes, for the most part, glacial meltwater remains at the surface due to the varying degrees of chemical stratification of the lakes, and there is little to no physical mixing within the lakes (SPIEGEL and PRISCU, 1998). The three lakes in Taylor Valley (Bonney, Hoare, and Fryxell), and Lake Vanda in Wright Valley, have substantially different climate histories and are situated in different landscape ages and positions, which have led to very different solute chemistries (LYONS et al., 2000). Lake Bonney, the westernmost lake, is separated into two lobes by the Bonney Riegel at a current depth of 14 m

(POREDA et al., 2004). The deeper portions of both lobes are saline to hypersaline. Lake Fryxell is approximately 6.5 km from the Ross Sea and is brackish below approximately 10 m. Lake Hoare is 10 km to the west of the Ross Sea and is the freshest of the lakes. Lake Vanda, a 72 m deep lake that has fresh surface waters and saline bottom waters. Like Taylor Valley lakes, Lake

Vanda has a perennial ice cover that forms a moat during the summer months. The chemistries and environmental histories of the saline lakes have recently been reviewed by Green and Lyons

(GREEN and LYONS, 2009).

2.4 Methods

Samples were collected during the 2005/06 and 2006/07 austral summers. Lake samples were collected by drilling a hole into the lake ice covers and retrieving samples using a Niskin bottle. Stream samples were collected using NalgeneTM polyethylene bottles. Lithium isotope and cation sample bottles were soaked in 10% HCl overnight and rinsed five times with 18 MΩ (DI)

58 water, whereas anion sample bottles were soaked overnight in DI water and rinsed three times with fresh DI water, before being sent to the field. Bottles were rinsed three times with stream water before being filled with sample. Samples were filtered through pre-cleaned filtering apparatuses using 0.4 µm Nucleopore filters within 12 hours of collection. Lake and stream cation samples were acidified to 2% v:v and 0.2% v:v, respectively, with trace-metal grade HNO3.

Snow samples were collected using the techniques of Witherow and others (WITHEROW et al.,

2006). Using standard avalanche shovels, snow pits were dug in the accumulation zones during the 2006-2007 field season. To eliminate the possibility of melting, the pit was oriented to keep the sampling face in a shadow during the entire sampling period. The sampling face was cleaned using a Teflon scraper. The Teflon scraper was “rinsed” by placing it in a snow layer adjacent to the one being sampled. Samples were placed in a one liter container, which was cleaned by soaking in DI water overnight and rinsing three times with DI water before being deployed to the field.

All aqueous samples were sent chilled to Crary Laboratory, McMurdo Station, Antarctica where they were analyzed for major ions on a Dionex DX-120 ion chromatograph (IC) following the protocols of Welch and others (WELCH et al., 1996). In order to determine the reproducibility of this analytical method, selected samples were run twice (Appendix B). Lithium concentrations were determined by an Elan Quadrupole ICPMS at the Trace Element Research Laboratory at

The Ohio State University. After the production of a calibration curve, a standard with a known concentration of lithium (referred to as a “check standard”) was analyzed to check the quality of the calibration and after every five samples to monitor instrument drift. Because the range in concentration in these samples is so large, glacier, stream, and lake samples were analyzed in separate batches. In addition to check standards, bottles filled with DI water (i.e. field blanks) were taken to the field and analyzed with the appropriate sample batch (Appendix B). Stream 59 sediments were collected by hand and stored in clean plastic containers both before and after air drying samples. Samples were analyzed by x-ray fluorescence spectroscopy at XRAL

Laboratories in Don Mills, Ontario, Canada.

Samples for lithium isotopic analysis were prepared and analyzed at the University of

Oxford. All methods for lithium isotope analyses were conducted in a metal-free ultra-clean laboratory and all who enter the laboratory are required to wear hooded, non-particulating

Tyvek™ suits, clean clogs, and powder-free latex gloves at all times. Briefly, the samples were dried and redissolved using a series of 16 M HNO3 and 0.1 M HNO3-50% methanol (MeOH)-DI solutions. The lithium in the samples was then separated using cation exchange. The 150 mL

Biorad columns contained 5 mL of AG50W-X8 resin which had been rinsed with HCl and pre- conditioned with a 0.1 M HNO3-50% MeOH-DI solution. The sample was extracted by a series

0.1 M HNO3-50% MeOH-DI and 1 M HNO3-80% MeOH-DI rinses. Lithium isotopes were analyzed by MC-ICPMS using modified techniques of Pistiner and Henderson (PISTINER and

HENDERSON, 2003) and Marriott and others (MARRIOTT et al., 2004) and replicate aliquots from selected samples were analyzed (Table 2.1). Lithium isotopic ratios are reported using the following equation:

7퐿푖 7퐿푖 − 6퐿푖 6퐿푖 푠푎푚푝푙푒 푠푡푎푛푑푎푟푑 훿 7퐿푖 = × 1000‰ 7퐿푖

6퐿푖 푠푡푎푛푑푎푟푑 relative to the L-SVEC Li standard. The average value of the LSVEC standard was 0.3‰ with a standard deviation of 0.4‰ (n=32).

60 Sample Name δ7Li δ7Li Duplicate

E. Lake Bonney 6 m 23.9‰ 26.8‰

Lake Fryxell 6 m 21.8‰ 28.6‰

Andersen Creek 19.5‰ 20.2‰

Lost Seal Stream 20.7‰ 21.0‰

Canada Glacier 2.8‰ 3.9‰

Table 2.1 Duplicate δ7Li Analysis.

2.5 Results

As indicated earlier, the lakes in the McMurdo Dry Valleys have varying degrees of salinity and chemical stratification due to differences in past and present environmental conditions (LYONS et al., 2005). This stratification persists due to the presence of perennial ice covers on the lakes which restrict physical mixing by wind. To illustrate the character of chemical stratification in the lakes, I present sodium concentration data as a function of depth

(Figure 2.1). Lake Hoare has the lowest surface water concentration which gradually increases with depth. The increase is most pronounced between 4 and 16 m. Lake Fryxell is more brackish than Lake Hoare and shows a linear increase (r2=0.99) in Na concentration with depth. The Na concentrations in the surface waters of Lake Bonney are similar to that of Lake Fryxell. From approximately 18-25 m in East Lake Bonney, the Na concentration rapidly increases by a factor of

2.7. In West Lake Bonney, the chemocline is at approximately 15 m where concentrations increase by 3.3 times. In Lake Vanda, the concentration of sodium above 55 m is relatively constant, but below 55 m, it drastically increases to over 6 times the concentration at 55 m. The

δ7Li values and the corresponding Li, Na, and Ca data for all of the samples analyzed are shown in Table 2.2. Taylor Valley glaciers have the lowest lithium concentrations and lightest δ7Li

61 values in the McMurdo Dry Valleys aquatic system. The mean concentration of lithium in glacier snow (0.17 µg/L) and the mean δ7Li value (+2.2‰) are well below the values found in the streams

(1.19 μg/L and +19.0 ‰) and lake surfaces (20.02 μg/L and +22.7‰) (Table 2.4). The average values reported above for the lake surface waters do not include Don Juan Pond, whose Li concentration is 2.65 x 105 µg/L (Table 2.4). In all lakes, water below the chemocline is isotopically heavier than water at the surface (Table 2.2, Figure 2.2)

A similar trend is seen in between the Li:Ca (mg:g) ratio and δ7Li values (Figure 2.3).

Commonwealth Glacier has a higher Li:Ca value than the other glaciers (Table 2.2). For all samples, the relationship between Li:Ca and δ7Li is strong and positive (r2 = 0.78). The hypersaline bottom waters of Lake Bonney do not follow this trend due to Li:Ca ratios that are at least 5-fold higher than the other lakes and streams. Li:Ca ratio and δ7Li values in Don Juan Pond surface water are very similar to the bottom water of West Bonney. The surface water of Lake

Fryxell also appears to not follow the same trend as most of the waters in the McMurdo Dry

Valleys.

When Li concentration data are normalized to Na concentrations, a strong, negative correlation (r2=0.84) emerges between Li:Na and δ7Li (Figure 2.4). Lake Vanda and Don Juan

Pond, the two lakes from Wright Valley, are the only exceptions. The surface water Li:Na ratio for Lake Vanda is about 20-fold higher than the average surface Li:Na ratio in Taylor

Na (mg/L) 0.1 1 10 100 1000 10000 100000 0

Depth (m) Depth 40

Hoare 50 Fryxell E. Bonney 60 W. Bonney Vanda Don Juan Pond 70

Figure 2.1 Sodium concentration by depth for the lakes examined in this study.

63 Depth Li Na Ca Li:Na Li:Ca δ7Li Sample Location (m) (µg/L) (mg/L) (mg/L) (mg:g) (mg:g) (‰)

Commonwealth Glacier 0.15 0.19 0.36 0.78 0.41 +0.8

Canada Glacier 0.10 0.20 2.52 0.49 0.04 +2.8 Glaciers Howard Glacier 0.25 0.41 3.33 0.61 0.08 +2.9

Andersen Creek 1.00 3.76 8.64 0.27 0.12 +19.5

McKnight Creek 1.50 11.66 8.04 0.13 0.19 +23.3

Wharton Creek 0.70 1.39 9.94 0.50 0.07 +11.6

Streams Priscu Stream 2.50 9.76 19.38 0.26 0.13 +18.0

Lost Seal Stream 0.80 4.15 4.01 0.19 0.20 +20.8

Onyx River 0.63 3.45 3.71 0.18 0.17 +20.6

5 3.0 32 18 0.09 0.17 +24.3 Lake Hoare 25 17.0 160 70 0.11 0.24 +25.9

6 46.0 296 59 0.16 0.79 +21.8 Lake Fryxell 15 95.0 2,160 125 0.04 0.76 +33.2

5 22.0 286 59 0.08 0.38 +17.3 W. Lake Bonney

35 4,400 39,200 2304 0.11 1.91 +23.5 Lakes 6 29.0 354 72 0.08 0.41 +20.1 E. Lake Bonney 35 8,290 48,500 1,640 0.17 5.04 +23.9

10 0.11 53 81 2.16 1.40 +30.3 Lake Vanda 70 22,000 5,000 21,000 4.41 1.05 +35.6

Don Juan Pond Surface 265,000 12,200 134,000 21.71 1.90 +22.1

Seawater 170 10,800 0.17 0.02 0.41 +30.8

64 +40

+35 Glaciers

+30 Streams

+25 Lake Hoare Li

7 +20

δ Lake Fryxell +15 Lake Bonney +10 Lake Vanda +5 Don Juan Pond +0 0.1 1 10 100 1000 10000 100000 1000000 Seawater

Li (μg/L)

+40

+35 Glaciers +30 Streams

+25 Lake Hoare Li

7 +20 Lake Fryxell δ Lake Bonney +15 Lake Vanda +10 Don Juan Pond +5 Seawater +0 0.01 0.1 1 10 Li:Ca (mg:g)

Figure 2.3 Lithium to calcium mass ratios versus δ7Li. Lighter symbols represent lake surface water samples, darker symbols represent bottom water samples, with the exceptions of Don Juan Pond and Seawater. Seawater data from (PILSON, 1998; ROSNER et al., 2007).

+40 Glaciers +35 Streams +30 Lake Hoare +25

Lake Fryxell Li

7 +20 δ Lake Bonney +15 Lake Vanda +10 Don Juan Pond +5 Seawater

+0 0.01 0.1 1 10 100 Li :Na (mg:g)

Sample Li:Na (mg:g) Location Sediments Water

Andersen Cr. 0.80 (n=2) 0.27

Lost Seal Str. 0.97 (n=3) 0.19

Priscu Str. 0.84 (n=2) 0.26

McKnight Cr. 0.55 (n=1) 0.13

Table 2.3 Average lithium and sodium composition of stream sediment samples compared to stream water samples.

Valley lakes. The Don Juan Pond water Li:Na ratio is 10-fold higher than that of Lake Vanda

(Table 2.2). Similarly, bottom samples from Lake Vanda shows Li:Na ratio that are about 40 times higher than the Li:Na ratio in bottom waters of the Taylor Valley lakes. The strong negative relationship between Li:Na and δ7Li is even more pronounced in the stream samples (r2

= 0.97, n=6). Of the stream sediments and water samples analyzed, Li:Na in the sediments are 2 to 5 times higher than in their respective streams (Table 2.3).

2.6 Discussion

2.6.1 Glaciers and aeolian input

The sources of solutes to the glacier surfaces are a combination of aeolian deposition of crustal material (aluminosilicates and simple salts), seawater components deposited as marine aerosols, and deposition from stratosphere aerosols rich in nitrogen (KEYS and WILLIAMS, 1981;

LYONS et al., 2003; MICHALSKI et al., 2005; WITHEROW et al., 2006). Dominant east-west winds are responsible for depositing carbonate and silicate dust onto glacier surfaces (FORTNER et al., 2005;

LYONS et al., 2003). Aeolian transport of dry valley salts is of particular importance with regards to mid-valley glaciers such as the Howard and Canada Glaciers which receive more sediment than the coastal glaciers (LANCASTER, 2002; WITHEROW et al., 2006). Coastal glaciers such as the

Commonwealth Glacier are more dependent on material delivered from the Ross Sea, especially during the late summer-autumn months when the sea ice is at a minimum (BERTLER et al., 2004b;

LEGRAND and MAYEWSKI, 1997; LYONS et al., 2003). Although the coastal glaciers are more dependent on marine inputs, debris from the valley floors contribute various amounts of dust to the all the glacier surfaces (LYONS et al., 2003; WITHEROW et al., 2006).

The δ7Li ratios of the glacial snow are amongst the lightest values observed to date in surface and groundwater samples (+0.8 to +2.9‰) and are very unlike seawater (+30.8‰)

67 (ROSNER et al., 2007; TOMASCAK, 2004). This suggests that marine aerosol or marine-derived salts are not the primary sources of lithium to the glacier surfaces, in contrast to previous assumptions

(LYONS and WELCH, 1997). Assuming all the Cl present in the snow is derived from marine aerosols and there is no fractionation between Li and Cl as the marine aerosols are forming, a simple calculation multiplying the ratio of Li:Cl in seawater (mg:g = 8.76 x 10-3) by the Cl concentration in Canada, Commonwealth, and Howard Glaciers (0.07 mg/L, 0.48 mg/L, and 11.5 mg/L) yields the contribution of Li from marine aerosols. Concentrations of 0.65 x 103 μg/L, 4.19 x

103 μg/L, and 11.5 x 103 μg/L for the Canada, Commonwealth and Howard Glaciers, respectively, indicate that marine aerosols account for only a very small fraction of Li to the glaciers (Table

2.2). Additionally, the Li:Na ratios of the glaciers (0.49—0.78 mg:g) is unlike seawater (0.02 mg:g) further demonstrating that marine-derived aerosols are not the primary source of Li to glacial snow.

Generally, geologic materials are isotopically lighter than natural water samples because of the preferential retention of 6Li in the solid phase. This depletion in solid substances is due to kinetic effects of fluid-particle interactions (TOMASCAK, 2004). Despite this general relationship,

Taylor Valley glacier snow samples have δ7Li signatures more like granite and rhyolite, which range from -1.2 to +8.0‰ (TOMASCAK, 2004). Although this wide range of Li isotopic values for granite does not conclusively mean the source to the glaciers is felsic rock, the snow samples are extremely light and resemble aluminosilicates. Thus, in accord with previous research, I conclude that the major source of Li in the glacier ice is derived from the dissolution of silicate materials or from Li containing salts (LYONS et al., 1998a).

68 2.6.2 Streams: Chemical weathering and salt dissolution

Lithium in solution is hydrated in the tetrahedral coordination state, but lithium in clays is in either the tetrahedral or octahedral sites (OLSHER et al., 1991). Incorporation of 6Li into octahedral sites is energetically preferred and can account for the observed distribution of heavier isotopes in solution compared to lighter isotopes in solids in aquatic systems. Lithium replaces Mg by coupled substitution in clays increasing the lithium concentration and 6Li (HUH et al., 1998). Pistiner and Henderson (PISTINER and HENDERSON, 2003) also observed preferential sorption of 6Li onto the octahedral sites of gibbsite. Such reactions not only occur within the stream channel, but they have been observed in the estuarine environment as well (POGGE VON

STRANDMANN et al., 2008). Inverse modeling of the major ion chemistry of the stream waters indicates that the 2:1 clay minerals, smectites, illites, and vermiculites may be produced due to silicate weathering in the stream channels (LYONS et al., 1998a). Furthermore, smectite has been observed in soils where moisture is present in particular, in undrained hollows where windblown sediment is entrapped or chemical weathering is occurring (CAMPBELL and

CLARIDGE, 1982).

Because of the preferential sorption of 6Li onto secondary clay minerals, Li isotopes can reflect the degree of chemical weathering in a catchment (HUH et al., 1998). In tectonically active watersheds, where the water:rock ratio is high, water initially leaches the heavier isotope from the bedrock resulting in high Li concentrations and heavy δ7Li values in stream water (HUH et al.,

2001). In the case of these high relief, reaction limited watersheds, chemical weathering is a unidirectional process as the products of chemical weathering are rapidly removed resulting in kinetic isotope fractionation, and equilibrium isotope fractionation most is prevalent in transport limited reactions (HUH et al., 2001). In transport limited watersheds with a low water to rock

69 ratio, the lithium concentrations in the water are lower than the bedrock, and the δ7Li is lighter than in reaction-limited watersheds (HUH et al., 2001). Furthermore, the suspended load has isotopic values that are lighter than the dissolved load regardless of weathering regime (HUH et al., 2001).

Lithium accumulates on the landscape due to long-term atmospheric deposition and the dissolution of these aerosols is at least a minor source of lithium to the aerosol-dominated streams (LYONS and WELCH, 1997). There are streams in the Fryxell Basin that have geochemistries dominated by chemical weathering and those dominated by marine salt dissolution (GREEN et al., 1988; GREEN et al., 1989; LYONS and WELCH, 1997). This is corroborated by the range of Li:Na values (LYONS and WELCH, 1997). It has been speculated that all streams in

Taylor Valley receive a large fraction of their lithium concentration from chemical weathering

(LYONS and WELCH, 1997). Chemical weathering in the hyporheic zones of dry valley streams contributes a significant amount of solutes particularly as one progresses downstream, thus longer streams have a higher solute load (GOOSEFF et al., 2002). With my limited stream data, I saw no relationship between δ7Li or lithium concentration and channel length, although Lyons and Welch (LYONS and WELCH, 1997) noted that longer streams (i.e. streams in the Lake Fryxell basin) usually had higher Li concentrations. These contrasting observations are likely due to the limited amount of data from streams in both datasets. It should be noted that the δ7Li values of the two streams in the Lake Fryxell Basin, Lost Seal Stream and McKnight Creek, are higher than the other streams in Taylor Valley which agrees with previous research on proglacial streams that shows increasing δ7Li values downstream from the glacier terminus (Table 2.2) (LEMARCHAND et al., 2008). This work supports the argument that chemical weathering can be a major source of Li to the streams.

70 Isotopic fractionation may also occur due to interactions with sediments in the streambed and in the hyporheic zone. Mass balance analysis of a lithium tracer experiment in Green Creek reveals that lithium is taken up and sodium is released from the stream-channel sediment

(GOOSEFF et al., 2004). It is postulated that in these reaches of the stream, lithium is exchanged for

Na, Ca, and/or Mg, or ion exchange and adsorption onto the negatively charged surfaces of sediments in the streambed and hyporheic zone (GOOSEFF et al., 2004). As noted above, the stream sediments in Taylor Valley are enriched in lithium relative to sodium compared to stream water. This may be an indication of lithium adsorption and ion exchange on the sediments, in such a case, 6Li would be preferentially removed from solution and adsorbed or exchanged on the sediments and further enriching the δ7Li value of the stream water.

2.6.3 Lakes and isotopic mixing

Previous research on lithium isotopes in terminal lakes has shown great isotopic variation (Table 2.4, Figure 2.5). Lithium isotopic values from Lake Tanganyika, the Caspian Sea, and the Dead Sea are remarkably similar to the mean seawater value (+32.0‰) (CHAN and

EDMOND, 1988). It is thought that these values are the result of fractionation associated with the incorporation of Li into evaporites or through “reverse weathering” reactions (CHAN and

EDMOND, 1988; VONDAMM and EDMOND, 1984). This suggests that the processes controlling the lithium isotopic composition in the ocean may also control the lithium isotopic composition in some large lakes such as these (TOMASCAK, 2004). Recent studies indicate that sorption onto clays rather than CaCO3 precipitation is the dominant mechanism for lithium removal in aqueous systems (ANGHEL et al., 2002; TOMASCAK et al., 2003). Sorption onto clays would result in the removal of 6Li from solution thereby enriching water in 7Li (CHAN and EDMOND, 1988). The lakes in the Great Basin are lighter than Lake Tanganyika, the Caspian Sea, and the Dead Sea and most

71 of the lakes in the McMurdo Dry Valleys (Figure 2.5). The isotopic values of inflow waters show that although clay sorption controls the δ7Li in some large lakes, mixing of streams and groundwater control the lithium isotopes in Great Basin Lakes (TOMASCAK et al., 2003). Lake samples collected from the Great Basin are consistently lighter than seawater due to the mixing of relatively light source water from streams (+11.3 to +29.4‰) and groundwater (+7.4 to +30.6‰)

(TOMASCAK et al., 2003). Furthermore, recent tufas in the Mono Lake basin are isotopically similar (+18.0‰) to the bulk of lake water (+20.1‰) demonstrating that CaCO3 precipitation has little effect on the isotopic composition of Mono Lake (TOMASCAK et al., 2003).

2.6.3.1 Lake Bonney

The western lobe of Lake Bonney is primarily fed by meltwater from the Taylor Glacier and Blood Falls. Blood Falls is a hypersaline brine flowing from under the Taylor Glacier with abundant amounts of halite, gypsum, aragonite, and Fe-oxyhydroxides, and it contributes significant amounts of solutes to West Lake Bonney (BLACK et al., 1965; LYONS et al., 1998b;

LYONS et al., 2005; LYONS et al., 1998c). Blood Falls has been interpreted to be chemically modified seawater isolated in Taylor Valley when the valley was a fjord during the Tertiary

(LYONS et al., 2005). Geologic evidence suggests that the sea regressed towards the east between

4.6 to 5.1 Ma and the Lake Bonney basin became isolated from its marine source (ELSTON and

BRESSLER, 1981; ISHMAN and RIECK, 1992). East Lake Bonney is primarily fed by water from West

Bonney, and a meltwater stream at the eastern end of the basin, Priscu Stream, contributes a small portion of water (SPIEGEL and PRISCU, 1998). The two lobes of Lake Bonney have different, yet related evolutionary histories that explain their geochemical differences (POREDA et al., 2004).

The western lobe of Lake Bonney has maintained a perennial ice cover and liquid water. East

Lake Bonney, on the other hand, was ice-free, lake levels were lowering due to evaporation, and

72 salts were precipitating out of the brine. Later, West Lake Bonney spilled over the Bonney Riegel, an ice-cover was formed on East Lake Bonney, and both lobes have continued to rise since approximately 200 yr BP. The relatively fresh water spilling from West Lake Bonney mixed very little with the brine below thus resulting in the chemical stratification of East Lake Bonney

(LYONS et al., 1998b).

Li (μg/L) δ7Li (‰)

Surface Bottom Surface Bottom Lake Hoare 3.0 17.0 +24.3 +25.9 This study

Lake Fryxell 46.0 95.0 +21.8 +33.2 This study

East Lake Bonney 29.3 8,290 +20.1 +23.9 This study

West Lake Bonney 22.4 4,400 +17.3 +23.5 This study

Lake Vanda 114 22,000 +30.3 +35.6 This study

Don Juan Pond 265,000 +22.1 This study

Mono Lake 11,000 +19.8 Tomascak et al., 2003

Great Salt Lake, UT 15,100 +16.7 Tomascak et al., 2003

Big Soda Lake, NV 625 +21.6 Tomascak et al., 2003

Walker Lake, NV 1,070 +23.3 Tomascak et al., 2003

Pyramid Lake, NV 720 +22.5 Tomascak et al., 2003

Lake Tanganyika 15 +33.3 Chan and Edmond, 1988

Caspian Sea 286 +32.1 Chan and Edmond, 1988

Dead Sea 13,700 +34.4 Chan and Edmond, 1988

73 +40

+35

+30

+25 McMurdo Dry Valleys

+20

Li (‰) Li 7

δ Great Basin +15 Other Terminal Lakes

+10 Seawater

+0 0.001 0.01 0.1 1 10 100 1000 1/Li (L/mg)

Because of the extreme solubility of lithium, these different evolutionary histories of the two lobes have had apparently little effect on the δ7Li values of the bottom waters. The monimolimnia of both lobes of Lake Bonney are isotopically similar indicating that the lithium in the bottoms of these lobes may have had the same source, and the East Bonney draw down had no effect on the Li isotopic composition. Earlier work examining the Li:Cl ratios in the monimolimnia excluded a hydrothermal source of lithium and speculated that the lithium was of marine origin (TAKAMATSU et al., 1998). The δ7Li values of the monimolimnia are lighter than seawater suggesting the contribution of lighter Li from chemical weathering in addition to the original seawater signature (Table 2.4). A similar argument has been made using other isotopic data (LYONS et al., 2002; LYONS et al., 2005). The Li:Na data demonstrate that the monimolimnetic waters of Lake Bonney are a mixture of seawater and glacial snow solutes (Figure 2.6). The 36Cl 74 data which show that the bottom waters of Lake Bonney lie on a seawater-glacial melt mixing line which may signify that conservative salts in the monimolimnetic waters of Lake Bonney are derived from the dilution of seawater by past glacial melt (LYONS et al., 1998c).

Lake Bonney Monimolimnia +35 SW R² = 0.98 90% +30 80% 70% +25 60% +20

50% Li 7 40% +15 δ 30% 20% +10 10% Glacial Snow +5

+0 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 Li:Na (mg:g)

The δ7Li values of the surface waters of both lobes of Lake Bonney are lighter than their monimolimnia and are lighter than expected if they were a two component mixture of seawater and solutes from precipitation. As noted above, west Lake Bonney is fed by glacier meltwater from Taylor Glacier, an outlet of the East Antarctic Ice Sheet. Because the origin of the ice at the current snout of Taylor Glacier, i.e. Taylor Dome is some 80 km away, its chemistry differs from the mountain glaciers in Taylor Valley. In contrast to the largely dust-dominated mountain glaciers in the Dry Valleys, the solutes in Taylor Dome ice are primarily of marine origin and may

75 have lithium isotope values quite different than the mountain glaciers in the dry valleys

(HINKLEY and MATSUMOTO, 2001). Additionally, the terminus of Taylor Glacier is submerged in

West Lake Bonney, and melt at this location flows directly into the lake. Because this water is not in contact with sediment, it inherits little lithium by chemical weathering or sea salt dissolution.

The ultimate source of Taylor Glacier melt water, limited chemical weathering interactions, and hypersaline contributions from Blood Falls may explain the lighter δ7Li values in the epilimnion of West Lake Bonney compared to other lakes in the region.

2.6.3.2 Lake Fryxell

Lake Fryxell has approximately 12 streams flowing into it that vary in length, discharge, chemistry, and morphology (CONOVITZ et al., 1998). Dissolved salts in the surface water of Lake

Fryxell are the result of present inflows of glacial meltwater modified by chemical weathering in the stream channel (GOOSEFF et al., 2002; GREEN et al., 1988; TAKAMATSU et al., 1998). This is also evidenced by the δ7Li values particularly when the lithium concentrations of the streams and the surface water of Lake Fryxell are normalized to Na (Figure 2.4).

Sedimentary materials obtained by deep drill holes in Taylor Valley next to the lake reveal that Lake Fryxell was influenced by intrusion of the Ross Sea at various times during the

Pliocene (PORTER and BEGET, 1981). The Ross Sea retreated between 1.7 and 1.8 Ma leaving behind a hypersaline playa (ELSTON and BRESSLER, 1981; ISHMAN and RIECK, 1992). This shallow pond, with a seawater-like composition, refilled approximately 1000 years ago (LYONS et al.,

1998b). The 36Cl/Cl, Li:Na, and halogen data all indicate that the solutes in the bottom of Lake

Fryxell are of marine origin (CARLSON et al., 1990; JONES and FAURE, 1978; LYONS et al., 2005;

TAKAMATSU et al., 1998). This is further corroborated by the δ7Li value of the bottom waters

(+33.2‰) compared to that of seawater (+30.8‰). It has been shown that during diffusion, 6Li is

76 more mobile than 7Li (FRITZ, 1992). Such diffusion of Li from the bottom brine in Lake Frxyell would result in an enrichment of δ7Li compared to seawater explaining the difference between the two values.

2.6.3.3 Lake Hoare

Lake Hoare is the freshest of the Taylor Valley lakes and its freshwater inputs are nearly exclusively from Andersen Creek and direct melt from the Canada Glacier. The lake was filled during the most recent advance of the Canada Glacier approximately 1000 years ago (LYONS et al., 1998b). Because of the young age of the lake, 36Cl, δ37Cl, and δ7Li values are very similar in the surface and bottom waters (LYONS et al., 1999b; LYONS et al., 1998a). This indicates that the solutes in the lake are derived from modern glacial melt. Unlike Lakes Bonney and Fryxell, any past desiccation events in Lake Hoare did not result in the extreme concentration of solutes

(LYONS et al., 1998b; LYONS et al., 1999b).

2.6.3.4 Lake Vanda

Lake Vanda is a meromictic lake separated into a saline layer below 54 m and a relatively fresh layer above that depth (ANGINO et al., 1965). The trace element chemistry of Lake Vanda is controlled by the “atmosphere-precipitation-glacier-glacial melt-stream-lake” pathway including crustal weathering in the stream channel (GREEN et al., 2005; MASUDA et al., 1982). The molar ratios of major ions in the Onyx River and water at 35 m are dissimilar indicating that the surface water is not entirely the result of concentration of river water (GREEN and CANFIELD, 1984). Much of the chloride, calcium, and lithium in the surface water of Lake Vanda (10 m) is diffused from the saline monimolimnion (GREEN and CANFIELD, 1984; TAKAMATSU et al., 1998). To calculate the

77 diffusional flux of lithium from the Onyx River and the bottom brine of Lake Vanda to the surface waters Fick’s equation:

푑푐 퐹 = −퐷 푑푥 was applied where F is the flux of ions, D is the diffusion coefficient, and dc/dx is the concentration gradient from 70 m to 10 m. Because the temperature increases with depth in Lake

Vanda, the diffusion coefficient used was 8.69 x 10-6 cm2/s at 18°C (LI and GREGORY, 1974).

Diffusion of Li from the saline brine is 0.10 mg/cm2/yr. The boundary between the brine and fresh water occurs at approximately 60 m and the cross-sectional area at this depth is 1.41 x 1010 cm2 (calculated from (CAMPBELL and WILSON, 1972)). Considering this cross sectional area, the brine contributes 1,410 kg/yr of lithium to the surface water, that is, water above 60 m. Since

1969, the annual discharge from the Onyx River ranged from 2.21 x 108 L to 1.72 x 1010 L

(MCKNIGHT, 2009). Multiplying these data with the concentration of Li in the river (0.63 μg/L) yields lithium inputs to Lake Vanda ranging from 0.13 to 10.8 kg/yr. Based on these calculations, diffusion of lithium from the deep brine is clearly the dominant source of lithium to the water above the chemocline.

Fritz (FRITZ, 1992) noted mass-dependent fractionation of aqueous LiCl during diffusion at 22°C where DLi6/DLi7= 1.011±0.0003. This ratio indicates that under a steady state, the light isotope becomes enriched on the positive pressure side, e.g. the hypersaline brine at the bottom of Lake Vanda. Because the boundary between this hypersaline brine and the fresher water above is not a membrane but rather a 10 m chemical gradient, this DLi6/DLi7 ratio would not produce appropriate results. Furthermore, the lithium isotopic data collected from Lake Vanda only represents a surface and bottom water sample. To better address the question of isotopic diffusion rates in the lake, more data is needed.

78 2.6.3.5 Don Juan Pond

Don Juan Pond is a hypersaline Ca-Cl brine fed primarily by groundwater with little surface water inputs (HARRIS and CARTWRIGHT, 1981). The pond is shallow, thus the hydrology is highly affected by the loss of water through evaporation. Evaporation is variable but is approximately 1-3 mm/day (HARRIS and CARTWRIGHT, 1981). The salinity of the pond is variable, and at times, it may be as high as 40 wt% salt (HARRIS and CARTWRIGHT, 1981). It is due to this high salinity, that it remains unfrozen throughout the year. The majority of the water in the pond comes from a dolerite aquifer, and like the pond itself, porewater surrounding the pond is a high- salinity Ca-Cl brine (TORII and YAMAGATA, 1981). Groundwater in the McMurdo Dry Valleys does not directly evolve from surface water, and it is highly unlikely that this Ca-Cl-rich groundwater originates from modern water (GREEN and CANFIELD, 1984; LYONS and MAYEWSKI,

1993). The results from this work show that although Don Juan Pond may have a δ7Li value similar to other waters in the McMurdo Dry Valleys, it is not the result of marine salts mixing with glacial meltwater (Figure 2.4).

Like Don Juan Pond, deep groundwater beneath the Precambrian Canadian Shield is a hypersaline Ca-Cl brine (FRAPE et al., 1984). It is thought that such deep brines may be formed when, during a glacial period, an advancing glacier overrides an inland sea (BEIN and ARAD,

1992). δ7Li values of these Canadian Shield brines are similar, yet slightly heavier, to seawater suggesting that brine solutes are marine in origin and that a small amount of Li has been lost to mineral uptake (BOTTOMLEY et al., 2003; BOTTOMLEY et al., 1999). The δ7Li value of Don Juan

Pond (+22.1‰) is also similar, although slightly lighter, than seawater. This suggests that if seawater was the source of the brine, lithium has not been lost to adsorption onto secondary minerals as such a process preferentially removes 6Li from solution. The lighter δ7Li value in Don

Juan Pond compared to seawater may be the result of Li uptake from the dolerite aquifer, as 7Li is 79 preferentially leached from the source rock. Furthermore, Li:Cl ratios indicate that Don Juan

Pond is extremely enriched (937.7 µg:g) in lithium compared to seawater (8.76 µg:g) implying an additional source of lithium.

2.7 Conclusions

By analyzing the lithium isotope chemistry of McMurdo Dry Valley waters, I have further identified the solute transport mechanisms in this polar ecosystem. My research shows that glaciers, streams, Lake Hoare, Fryxell, and Bonney construct a mixing curve between primary precipitation and seawater. This is especially well illustrated because of the marine- derived bottom waters of Lake Fryxell. The monimolimnetic waters of Lake Bonney are derived from past dilution of seawater by glacial melt. The lithium in the surface waters of Lake Vanda are the result of Onyx River water and the diffusion of solutes from saline bottom waters. Don

Juan Pond is not characterized by evapoconcentration of glacial melt; rather its water source is most likely groundwater.

2.8 Acknowledgements

I am grateful for the following people for their assistance: Dr. Gideon Henderson and

Andrew Mason for help with analysis of lithium isotopes; Dr. Carolyn Dowling and Jill Gudding for collecting stream sediment samples; Kathleen Welch for major element analysis; Dr. John

Olesik and Anthony Lutton for their assistance with lithium concentration analysis; the Montana

State University Limno Team for sample collection; Drs. W. Berry Lyons, Anne Carey, Ozeas

Costa, Bill Green, and Dave Porinchu for their thoughtful comments on the manuscript; Chris

Gardner for providing the map for this paper. This work was supported by NSF-ANT grant

80 #0423595, a GSA Graduate Student Research Grant, and an OSU Office of International Affairs

Travel Grant.

Chapter 3 The Fate of Minor Alkali Elements in the Chemical Evolution of Closed-Basin Lakes

3.1 Abstract

The ultimate chemistry of a closed basin lake is determined at a very early stage by the interaction of dilute inflow waters with the surrounding geology. Alkaline earth elements and alkali metals play the most important role in these processes as the final brine type is defined by the abundance of these elements. The role of major ions in brine evolution has been studied in great detail, but little has been done on the behavior of minor alkali elements in these systems despite their similar chemical affinities to the major cations Na, K, Mg, and Ca. I have examined three major anionic brine types, chloride, sulfate, and bicarbonate-carbonate, in fifteen lakes in

North America and Antarctica to determine the geochemical behavior of lithium, rubidium, strontium, and barium in each brine type. Lithium and rubidium are largely conservative in all water types, and their concentrations are the result of long-term solute input and concentration through evaporation and/or sublimation. Strontium and barium behaviors vary with anionic brine type. Strontium can be removed in sulfate and carbonate-rich lakes by the precipitation of carbonate minerals. Barium may be removed in chloride and sulfate brines by either the precipitation of barite or perhaps biological uptake.

3.2 Introduction

The hydrologic balance of many arid and semi-arid lakes is determined by the input of water through springs, streams, and very limited precipitation and the loss of water through evaporation or sublimation. Such a balance can result in the formation of a saline lake if the 82 outflow of water is restricted (i.e., it is hydrologically closed), the evaporation exceeds inflow, and the inflow is sufficient to provide a standing body of water (EUGSTER and HARDIE, 1978).

Given these conditions, the ultimate chemistry of a closed-basin lake is determined by the initial composition of precipitation, the weathering reactions between dilute inflow water and lithology, and evapoconcentration (or sublimation) of the lake water (HARDIE and EUGSTER, 1970).

The chemical pathways of evolving brines are determined very early by the initial chemical composition. These pathways have been termed “geochemical divides” (Figure 3.1).

That is, during mineral precipitation from a brine, the less abundant ion of the mineral pair of ions will become drastically depleted compared to the other (HARDIE and EUGSTER, 1970). These models assume that once solids reach saturation in a brine, they are removed and no longer interact with the water. Because of its lower solubility, calcium carbonate (CaCO3) precipitation is the first geochemical divide encountered during brine evolution. Despite the brine type,

CaCO3 is always to first salt to reach saturation and distinguish carbonate-rich from carbonate- poor brines (HARDIE and EUGSTER, 1970). It is at this point where the lake becomes either enriched in HCO3+CO3 or Ca+Mg depending on the (HCO3+CO3)/(Ca+Mg) ratio of the inflow.

Alkaline earth elements may become further enriched or depleted in these waters due to coprecipitation of other minerals such as sepiolite (Mg4Si6O15(OH)2·6H2O) (HARDIE and EUGSTER,

1970). Subsequent evaporite minerals such as gypsum and mirabilite can be critical in the formation of other geochemical divides and the enrichment in ions such as Ca, Mg, SO4, and Na

(HARDIE and EUGSTER, 1970).

Figure 3.1 The geochemical evolution of closed-basin lakes as determined by major ion concentration of waters and minerals precipitated (HARDIE and EUGSTER, 1970).

Early modeling of closed-basin lakes focused on the major elements of most natural waters: Ca, Na, K, Mg, Cl, SO4, and carbonate alkalinity (HCO3 + CO3) and the simple salts that these ions produced during evaporation (EUGSTER and JONES, 1979; HARDIE and EUGSTER, 1970).

This research was fundamental in illustrating how the geochemistry of a lake can result from dilute inflow water with a composition unlike the final brine. It is clear from this model, the abundance of major alkali elements, Ca, Na, K, and Mg, is key in determining the geochemical pathways involved in brine formation, yet there has been little interest in extending the modeling efforts to aid in the prediction of minor and trace metal behavior during brine formation. In this paper, I have made an attempt to understand minor alkali metal and alkaline earth behavior in different brine types in order to extend the earlier efforts on major cations. I have collected and analyzed water samples from brackish and saline lakes in the Great Basin of the United States,

Saskatchewan, and the McMurdo Dry Valleys of Antarctica (Figure 3.2, Appendix A, Appendix

E, Appendix F). The dataset includes surface samples from selected sulfate and carbonate brines and complete geochemical profiles of four chloride-type lakes. These data have been manipulated using the thermodynamic program, PHREEQ, to examine simultaneous equilibrium reactions in the lakes. These thermodynamic reactions at various salinities and chemistries illustrate the potential removal mechanisms that are occurring in closed-basin lakes.

3.3 Site Description

3.3.1 McMurdo Dry Valleys

In southern Victoria Land, Antarctica, the Trans-Antarctic Mountains block ice flow from the East Antarctic Ice Sheet towards McMurdo Sound. This results in the largest ice-free area in

Antarctica (approximately 4800 km2) known as the McMurdo Dry Valleys (MCM) (Appendix A)

(CHINN, 1990). The MCM (≈78° S) is a polar desert classified by extremely low temperatures and

Figure 3.2 The locations of the sites examined in this study.

high aridity. The mean annual temperature ranges from –16° C to –21° C, and it receives less than 10 cm of precipitation per year in the form of snowfall (FOUNTAIN et al., 1999b). Snowfall on the valley floor typically sublimates before melting and does not contribute significant amounts of water to the streams and lakes (CHINN, 1981). Even in these harsh conditions, perennially ice- covered lakes and hypersaline ponds exist in Taylor Valley (Lakes Fryxell, Hoare, and Bonney) and in Wright Valley (Lake Vanda and Don Juan Pond) (CHINN, 1993).

The 24 ephemeral streams in the MCM flow between 6-10 weeks during the austral summer from November to January and flow is highly variable both daily and seasonally

(CONOVITZ et al., 1998; FOUNTAIN et al., 1999b). The water in these ephemeral streams is derived from glacial melt and it is the sole source of water to closed-basin lakes of the MCM. Like temperate terminal lakes, MCM lake levels are sensitive to small changes in water inflow (CHINN,

1993; DORAN et al., 2008). The only means of maintaining hydrological balance in the lakes is by

86 sublimation of the perennial ice covers, which are replenished by the freezing of lake water on the bottom of the ice (CLOW et al., 1988). During cold periods, sublimation is greater than stream input, thereby lowering lake levels. Between 1903 and 1990, lake levels rose at an average rate of

16 cm/yr at Lake Bonney (CHINN, 1993). Since 1990, generally, lake levels have been decreasing in response to cooler summer temperatures (DORAN et al., 2002; DORAN et al., 2008). The location of lakes in the valley coupled with how they respond to long-term climatic fluctuations have resulted in salinities ranging from fresh to hypersaline (LYONS et al., 1998a).

Although the salinities between the lakes vary greatly (TDS = 40-645,000 mg/L), all of the

MCM lakes have Cl as the major anion (GREEN et al., 1988; LYONS et al., 1998a). The glacial melt inflow is enriched in marine aerosols, and the dissolution of aeolian deposited marine salts contributes to enrichment in chloride (LYONS et al., 1998a; WITHEROW et al., 2006). Sodium is the dominant cation in the Taylor Valley Lakes (Bonney, Fryxell, and Hoare) reflecting the influence of marine salts on stream chemistry and all these lakes can be classified as Na-Cl type waters.

The hypolimnion of Lake Vanda and Don Juan Pond, both in Wright Valley, are unlike the Taylor

Valley Lakes in that they are Ca-Cl brines. The sources of solutes in these lakes differ from those in Taylor Valley. The ultimate source of these solutes in the hypolimnion of Lake Vanda has been highly debated and may be the result of a previous drawdown of the lake, groundwater inputs, or relict seawater (ANGINO et al., 1965; LYONS et al., 2005). The lakes with the highest relative percent of chloride are the more geochemically evolved hypersaline lakes: Lake Bonney, Lake

Vanda, and Don Juan Pond. It is in the hypolimnia of Lakes Bonney and Vanda and in Don Juan

Pond that much of the initially dissolved bicarbonate and sulfate have been removed as CaCO3 and gypsum (GREEN and LYONS, 2009). Because of the conservative nature of chloride, it stays in solution and is concentrated during evaporation.

87 3.3.2 Great Basin

3.3.2.1 Pleistocene Lake Lahontan remnants

The geochemistry of Great Basin lakes has been studied in great detail since the 1960s and was key in developing the brine evolution model (JONES, 1966). These terminal lakes in the western United States are remnants of large pluvial lakes that occupied the region during wetter periods in the Pleistocene (Appendix E) (BENSON et al., 2002a). Mono Lake, Pyramid Lake, and

Walker Lake, are the remnants of Pleistocene Lake Lahontan, and are characterized by relatively high concentrations of carbonate and bicarbonate. The Lake Lahontan basin waters actively precipitate carbonate minerals until all alkaline earths have been depleted (JONES, 1966). As the lakes evaporated, calcium and magnesium are removed as sulfate and carbonate minerals.

Reducing conditions may also play an important role in removing the sulfate from the lakes particularly those that are more evolved (EUGSTER and HARDIE, 1978).

Pyramid Lake, located approximately 50 km northeast of Reno, Nevada, is fed by the

Truckee River system which drains Lake Tahoe and a number of other smaller lakes (BENSON and

WHITE, 1994). Since 1906, more than 50% of the water in the Truckee River has been diverted to the Lahontan Reservoir, and after 1938, Pyramid Lake became hydrologically closed (BENSON,

1994). Stream discharge has been reduced by 60% since the late 19th century as a result of mining and irrigation needs (BENSON and LEACH, 1979). Since the late 1860s, the salinity (TDS) of the lake has risen from approximately 3,300 mg/L to over 5,500 mg/L (GALAT and JACOBSON, 1985;

GALAT et al., 1981). The major ions in the lake are predominately Na, Cl, and HCO3+CO3 (GALAT et al., 1981). Periodic whiting events dominated by aragonite have been ascribed to increased allochthonous inputs of calcium, dissolution of calcite in lake sediments, water temperature, and biogenic factors (GALAT and JACOBSON, 1985).

88 Like many endorheic basins, stream water in the Walker basin have been diverted from the lake since European settlement. These diversions have resulted in a lake surface level decline of more than 45 m and, total dissolved solids have increased from 2,500 to 16,000 mg/L since 1882

(LOPES and SMITH, 2007). Walker Lake is a monomictic, moderately saline brine (TDS = 12,000-

13,000 mg/L) with Na as its primary cation and CO3 as its major anion (BUETEL et al., 2001;

DOMAGALSKI et al., 1989). During the summers, water temperatures rise from 6° C to 22.5° C in the surface water and to 14° C in the bottom water fully stratifying the lake by the early summer

(BENSON and SPENCER, 1983). By late summer, CaCO3 precipitates due to the warm temperatures

(BENSON et al., 1991; GALAT and JACOBSON, 1985). Incoming fresh water from the Walker River is responsible for the winter overturning of the thermally stratified lake (BENSON and PAILLET,

2002b). Despite the reduction in inflow to the lake, it is at an early stage of chemical evolution

(DOMAGALSKI et al., 1989).

Mono Lake lies in a tectonic basin in Eastern California (38° N, 119° W, 1,943 masl)

(JELLISON and MELACK, 1993). The lake covers 160 km2 and is meromictic (JELLISON and MELACK,

1993). It is fed by six perennial streams and Ca-Na-HCO3-rich springs (BISCHOFF et al., 1993b).

Since 1941, the largest two streams have been diverted to Los Angeles as a municipal water supply, which has resulted in a 50% reduction in lake volume (BENSON et al., 2003). The saline

(TDS = 97,000 mg/L) brine is dominated by CO3 and Na as its major ions (BISCHOFF et al., 1993b).

The lake is at an advanced state of evolution and is presently precipitating gaylussite

(Na2Ca(CO3)2 · 5H2O) and ikiate (CaCO3 · 6H2O) in addition to tufa towers of aragonite, calcite, and amorphous CaCO3 (BISCHOFF et al., 1993b).

89 3.3.2.2 Pleistocene Chewaucan Lake remnants

Lake Abert and Goose Lake are remnants of Pleistocene Lake Chewaucan in southern

Oregon and on the border between California and Oregon, respectively (Appendix F). Mean annual temperatures in the northern reaches of the Great Basin range from -23° to 32° C (PHILLIPS and VAN DENBURGH, 1971). Most of the water entering Lake Abert and Goose Lake is from snowmelt at higher elevations in the surrounding mountains (300-1500 masl) (PHILLIPS and VAN

DENBURGH, 1971). The primary inflow to Lake Abert is the Chewaucan River, and Goose Lake receives its water from numerous smaller streams. Additional sources include direct precipitation on the lakes and peripheral springs. Both of the lakes have undergone periods of total evaporation, and Goose Lake overflowed between 1868 and 1891 (PHILLIPS and VAN

DENBURGH, 1971). During the 20th and 21st centuries, inflow water has been diverted for agriculture, so future overflow is unlikely (PHILLIPS and VAN DENBURGH, 1971). Lake Abert is a

Na-Cl-CO3 brine with total dissolved solids ranging from 18,700-95,000 mg/L (CARLSON et al.,

1990). Because it has been periodically refreshed by overflow, Goose Lake, a Na-HCO3 brine, is much fresher with total dissolved solids (TDS) ranging from 600 to 2,700 mg/L (CARLSON et al.,

1990). The wide range in salinities in these lakes is due to climatic conditions and anthropogenic activities which dictate the balance of evaporation and inflow. In Abert and Goose Lakes, it has been suggested that carbonates and silicates are being removed by inorganic precipitation and in the case of silicates, biogenic processes (PHILLIPS and VAN DENBURGH, 1971).

3.3.2.3 Canadian lakes

It is estimated that the Great Plains of Canada have between one and ten million saline lakes (Appendix F) (LAST and GINN, 2005). The region is relatively flat with Pleistocene till and drift overlying Cenozoic and Tertiary bedrock (LAST, 1989). The lakes are young (<10,000 years

90 old), and most are remnants of larger extinct proglacial lakes such as Lake Regina, Lake Hind,

Lake Saskatchewan, and Lake Agassiz (LAST and GINN, 2005). Like many saline lakes, water levels, salinity, and organic productivity in northern Great Plains lakes are sensitive to climatic changes (LAST and SLEZAK, 1988).

The area is characterized as a cold, semi-arid climate, and although the mean annual temperature is 3° C, the low humidity, high winds, and warm summer temperatures are conducive to high evaporation rates (LAST, 1989). These rapid evaporation rates coupled with large interior drainage basins are the most important physical conditions in the formation of these lakes (LAST, 1989). During the cold, clear winter months, many, but not all, lakes form an ice cover (LAST and GINN, 2005). Many of the lakes in the region are shallow (<3 m) and demonstrate playa characteristics, that is, they go through periods of drying and refilling (LAST and SCHWEYEN, 1983).

The lakes show a wide range in salinity from brackish (TDS = 1,000-5,000 mg/L) to saline

(TDS > 5,000 mg/L), and the chemistry is primarily dominated by sulfate (LAST and GINN, 2005).

There are spatial and regional setting patterns influencing lake chemistry (LAST and GINN, 2005).

Sulfate along with Mg, Ca, and Cl demonstrate considerable range and relative concentrations between lakes (LAST, 1989). In western Saskatchewan and eastern Alberta, the shallow groundwater is somewhat fresh (TDS < 3,000 mg/L) and dominated by SO4 (LAST, 1992). It is clear that groundwater input plays a key role in the geochemistry of the lakes (BETCHER et al.,

1995).

The lakes in Saskatchewan were chosen because of their relatively high sulfate concentrations. Ceylon Lake, Freefight Lake, and Waldsea Lake are dominated by sulfate and

Deadmoose Lake and Little Manitou Lake are mixed SO4-Cl brines. In all of the lakes, sodium is the dominant cation indicating that either the inflow waters were highly depleted in calcium and 91 magnesium and/or that these cations have been removed by the precipitation salts. The larger, perennial lakes are chemically and thermally stratified with complex mineral precipitation/dissolution profiles (LAST and SCHWEYEN, 1983). The lakes precipitate calcite and high magnesium calcite and periodic whiting has been observed in Deadmoose and Waldsea

Lakes (LAST, 1989). The closed basin lakes of Saskatchewan are strongly influenced by annual and daily fluctuations in temperature. For example, some lakes become supersaturated with respect to a mineral when winter temperatures drop and ice covers form increasing salinity in the surface waters (LAST and GINN, 2005). These minerals fall to the undersaturated depths of the mixolimnion where the endothermic reaction of dissolution further cools the surrounding water

(LAST and GINN, 2005). This results in sedimentation of calcium and calcium-magnesium carbonates and sodium, magnesium, and sodium-magnesium sulfates (LAST and GINN, 2005).

All of these lakes were selected because they are either brackish or saline or, in the case of

MCM lakes, their hypolimnia are brackish or saline. Because their major anion concentrations vary greatly, the comparison of these very different closed-basin systems provide an excellent test of the importance of anionic composition on minor-element mineral precipitation.

3.4 Methods

3.4.1 Sample collection

Sample collection bottles for cations were cleaned by soaking the bottles in dilute HCl and rinsing them with 18 MΩ water (DI). Anion collection bottles were soaked and rinsed in DI.

Stream samples were collected in pre-cleaned Nalgene polyethylene bottles after rinsing the bottles with sample water three times. For the Great Basin and Canadian lakes, only surface samples were collected while depth profiles are available for the Antarctic lakes. Samples were collected from the McMurdo Dry Valley lakes at various depths by lowering a Niskin bottle into

92 a hole drilled into the ice and filling a pre-cleaned Nalgene polyethylene bottle (LYONS et al.,

1998a). Great Basin lake samples were collected in pre-cleaned polyethylene bottles. After collection, samples were stored in the dark and sent to The Ohio State University. All samples from Antarctica and the Great Basin were filtered through 0.4 µm Nuclepore filters and cation samples were preserved with trace-metal grade HNO3 within 1 week after collection. Samples from Saskatchewan were obtained from the archives at the University of Manitoba (LAST, 1992).

These samples were filtered through 0.4 μm Nuclepore filters and preserved with trace-metal grade HNO3 at The Ohio State University.

3.4.2 Laboratory analysis

Samples and field blanks were analyzed for Cl, SO4, Na, K, Mg, and Ca on Dionex DX-

120 Ion Chromotograph (IC) using modified techniques from (WELCH et al., 1996). An independent multi-element standard was analyzed after creating a calibration curve to satisfy quality assurance and quality control. In order to determine the reproducibility of this analytical method, selected samples were analyzed twice (Appendix H).

Li, Rb, Sr, and Ba were quantified on a Perkin-Elmer Sciex ELAN 6000 Inductively

Coupled Plasma Mass spectrometer (ICPMS). An attempt was made to quantify Cs concentrations, however all sample were below the detection limit of 0.05 μg/L. In addition to sample analysis, the concentrations of Li, Rb, Sr, and Ba were quantified in field blanks (i.e. bottles filled with DI that were transported to the field and processed as samples) (Appendix H).

Multi-element standards were used for six and eight point calibration curves. Calibration curves were analyzed at least twice in an analytical run. For approximately half of the samples, instrument drift was monitored by using a mid-concentration check standard every 1 to 5 samples. The reproducibility of this standard is illustrated by the relative standard deviations in

93 Appendix H. For the other samples, a 20 μg/L internal standard of Be, Co, La, and Y was implemented for the high salinity samples as a means of allowing the instrument to correct for drift.

Sample aliquots from the McMurdo Dry Valley lakes were analyzed for dissolved inorganic carbon (DIC). The samples were bottled with no head space and preserved with 0.5% v:v chloroform. The samples were stored chilled in the dark until analysis by either a Lira IRGA with an HP integrator or a Licor 6252 system in the Crary Lab at McMurdo Station, Antarctica

(MCM-LTER, 2007). Because the lakes are stratified, DIC replicate measurements below the chemoclines can be compared to illustrate the precision of the method (Appendix B). Alkalinity for the Great Basin and Canadian lakes was analyzed at The Ohio State University. This was achieved by titratrating 10 mL sample aliquots with 1 N Optima HCl. The precision of this method has been shown to be within 5% (STANDARD METHODS, 1999).

3.4.3 Geochemical modeling

Saturation indices were analyzed by PHREEQCI, a graphical interface of the thermodynamic model PHREEQ (CHARLTON and PARKHURST, 2002; PARKHURST and APPELO,

1999). Great Basin and Canadian lake indices were calculated at 20° C and 0° C, and MCM lake indices were calculated based on in-situ temperature measurements. PHREEQ is a general geochemical modeling program that is suitable for many different aqueous environments, but it is necessary to illustrate some of the program’s limitations. The default program uses ion associations and Debye-Hückle equations to determine the saturation indices of a suite of minerals. Due to the high ionic strength of the solutions in this study, Debye-Hückle equations are not appropriate thus a modified Pitzer database was used. The modifications were the addition of thermodynamic data for reactions and species not included in the Pitzer database

94 (Table 3.1) (PARKHURST and APPELO, 1999). It should be noted that the initial database was a compilation of equilibrium constants and enthalpies from various sources (PARKHURST and

APPELO, 1999). The thermodynamic information used to modify the Pitzer database was gleaned from the Minteq database included in the PHREEQCI package (CHARLTON and PARKHURST, 2002).

Barite was not included in the Minteq database, so its thermodynamic data were tabulated from

Stumm and Morgan (1996).

3.5 Results

3.1.1. Concentrations: Major elements

The salinities, as total dissolved solids (TDS), of the Canadian lakes ranged from 2.7 to

23%. (Appendix G). The highest TDS values were in Ceylon Lake, and the lowest were in

Deadmoose Lake. The salinities of the lakes in the Great Basin ranged from 0.4 (Pyramid Lake) to

12% (Abert Lake). The range of the salinity of the dry valley lakes was 0.003 to 44% with the lowest values in the surface water of Lake Hoare and the highest salinity in Don Juan Pond.

Although salinities ranged from very fresh to hypersaline, some generalizations can be made regarding the chemical character of the water (Figure 3.3).

The pH in the MCM lake profiles ranged from 5.74 in the hypolimnion of W. Lake

Bonney to 9.05 in the epilimnion of Lake Hoare. Sodium was the dominant cation in nearly all of the lakes in the MCM with the exception of Lake Vanda and Don Juan Pond, where Ca was dominant. The surface water of Lake Hoare has the lowest concentrations of Na (20.3 mg/L), K

(3.6 mg/L), Mg (3.64 mg/L), and Ca (13 mg/L). The water at 35 m in E. Lake Bonney had the highest concentrations of Na (48,500 mg/L), K (2,780 mg/L), and Mg (31,500 mg/L). Don Juan

Pond had the highest concentration of Ca (140,000 mg/L). All of the McMurdo Dry Valley lakes/ponds had chloride as the dominant anion. The water at 4 m in Lake Hoare had the lowest

Reactions: Products: Phases:

H4SiO4 = H4SiO4 SrOH Strontianite: SrCO3

H4SiO4 = H3SiO4- + H+ SrHCO3 Witherite: BaCO3

H4SiO4 = H2SiO42- + 2H+ SrCO3 Amorphous Silica: SiO2(a)

CO32- + 10H+ + 8e- = CH4 + H2O SrSO4 Calcedony: SiO2

Ca2+ + H2O = CaOH + H+ LiOH Quartz: SiO2

Ca2+ + CO32- + H+ = CaHCO3+ LiSO4 Sepiolite (d): Mg4Si6O15(OH)2·6(H2O)

Ca2+ + HSO4- = CaHSO4+ SiF62- Sepiolite: Mg4Si6O15(OH)2·6(H2O)

Ba2+ + H2O = BaOH- + H+ MgF+ BaF2

Ba2+ + CO32- = BaCO3 CaF+ SrF2

Ba2+ + HCO3- = BaHCO3+ NaF Barite: BaSO4

Ba2+ + SO42- = BaSO4 HF

Rb+ = Rb+ HF2-

F- = F- H2F2

Table 3.1 Reactions, products and phases added to the Pitzer database in PHREEQ.

Na+K Ca Cl GBGreat Basin

MCMMcMurdo Dry Valleys SKCanada

Mg SO4 CO3+HCO3

Figure 3.3 Piper diagram of the lakes in this study.

concentrations of Cl (33.5 mg/L) and SO4 (13.5 mg/L). The surface water of Lake Hoare and Lake

Vanda had the lowest concentration of HCO3 + CO3 (10.3 mgC/L). Water from Don Juan Pond was not analyzed for HCO3 + CO3 as calculating the charge balance of the water indicated the alkalinity would only constitute 0.26% (meq) of the total dissolved solids. The water at 37 m in

W. Lake Bonney had the highest HCO3 + CO3 (770.5 mgC/L) and SO4 (4,600 mg/L) concentrations.

Don Juan Pond had the highest Cl concentrations of all lakes in the MCM (283,000 mg/L).

The pH in the Saskatchewan Lakes (surface samples) ranged from 7.88 in Ceylon Lake to

8.54 in Waldsea Lake. Na was the dominant cation in the lakes with concentrations ranging from

2,160 mg/L in Waldsea Lake to 196,000 mg/L in Ceylon Lake. Deadmoose Lake had the lowest concentrations of K (420 mg/L) and Mg (525 mg/L), and Ceylon Lake had the highest concentrations of these cations, 2,600 mg/L and 2,320, respectively. The highest concentration of

Ca was in Waldsea Lake (120 mg/L), and the lowest was in Freefight Lake (7.9 mg/L). These lakes

97 in Saskatchewan had high concentrations of sulfate, and in the cases of Little Manitou Lake and

Deadmoose Lake, chloride was also abundant. Sulfate concentrations ranged from 14,800 mg/L in Deadmoose Lake to 196,000 mg/L in Ceylon Lake. The highest concentration of Cl was in

Little Manitou Lake (21,300 mg/L), and the lowest concentration occured in Freefight Lake (4,190 mg/L). HCO3 + CO3 concentrations ranged from 728 mgC/L in Deadmoose Lake to 5,958 mgC/L in Freefight Lake.

The lakes in the Great Basin had Na as the dominant cation, and pHs were alkaline ranging from 9.22 in Pyramid Lake to 9.88 in Abert Lake. Mono Lake had the highest concentrations of K and Ca, 1,500 and 3.7 mg/L, respectively. Magnesium and Ca were below the detection limit in Abert Lake; this has been shown in previous work as well (EUGSTER and

HARDIE, 1978). Pyramid Lake has the lowest concentration of K (94 mg/L) and Walker Lake had the highest concentration of Mg (40.3 mg/L). Mono and Goose Lakes were carbonate-rich,

Pyramid Lake was chloride-rich, and Abert and Walker Lakes were enriched in both HCO3+CO3 and Cl. HCO3 + CO3 concentrations ranged from 1.330 mgC/L in Pyramid Lake to 36,700 mgC/L in Abert Lake. Chloride ranged from 602 mg/L in Goose Lake to 29,800 mgC/L in Abert Lake.

Sulfate ranged from 276 mg/L in Pyramid Lake to 8,970 mg/L in Mono Lake. The major ion concentrations in Abert, Pyramid, and Mono Lakes have been previously studied in great detail

(Error! Reference source not found.). Comparison of the data generated in this study to that published by Eugster and Hardie (1978), shows that the solutes above the detection limit in Abert

Lake have decreased dramatically since the mid 20th century. Most solutes in Pyramid Lake water have also decreased over the past several decades, the exceptions being SO4 and

HCO3+CO3, which showed little change since reported in 1978. Mono Lake water has increased in Cl, SO4, Na, K, and HCO3+CO3 concentrations, and Mg and Ca concentrations have decreased since 1978. 98 The pH values in the MCM lakes were generally lower than the lakes in Saskatchewan, which were lower than the alkaline lakes of the Great Basin. Of all the lakes studied, the Great

Basin lakes had the highest concentrations of HCO3 + CO3, and the Saskatchewan lakes had the highest concentrations of SO4. The hypersaline waters of the MCM had the highest Cl, Ca, and

Mg concentrations and the surface waters had the lowest. So, I shall term the Great Basin lakes, the “carbonate lakes,” the Saskatchewan lakes, the “sulfate lakes,” and the McMurdo Dry Valley lakes, the “chloride lakes.”

3.5.1 Concentrations: Minor elements

Of the minor elements, Li, Rb, Sr, and Ba, all of the Canadian and Great Basin lakes had

Li in the highest concentration except Walker Lake which had Sr concentrations higher than Li.

Lake Hoare, Lake Fryxell, and West Lake Bonney were dominated by high concentrations of Sr.

Lake Vanda, Don Juan Pond, and the monimolimnion of East Lake Bonney had high concentrations of Li and Sr (Appendix D). Lithium and Rb demonstrated positive relationships with Cl (Figure 3.4, Figure 3.5). Strontium had a strong relationship with calcium but a poor relationship with chloride (Figure 3.7, Figure 3.8). Barium did not correlate well with any major or minor elements including Cl (Figure 3.6).

The Canadian lakes had the highest Li:Cl values which ranged from 230 to 1,900 mg:g.

The Great Basin lake water ranged from 2 to 560 Li:Cl (mg:g). The Li:Cl value in Don Juan Pond water was 940 mg:g, and the other lakes in the MCM ranged from 27 to 450 mg:g. The Rb:Cl value in Mono Lake water was much higher than all other lakes in this study, 100 mg:g. The

Rb:Cl in other GB lakes ranged from 9 to 19 mg:g. The Saskatchewan lakes discussed in this paper ranged from 2-17 Rb:Cl (mg:g). Don Juan Pond had the lowest Rb:Cl value of 0.4 mg:g, which was lower that the values in other McMurdo Dry Valley lakes (1.4-63 mg:g).

99 Because the MCM lakes were sampled at depth, it is possible to examine the Li:Cl values throughout the water column. Generally speaking, Lakes Bonney, Vanda, and Hoare maintained a constant ratio with depth, but the Li:Cl values in Lake Fryxell gradually decreased with increasing depth. Rb:Cl values strongly decreased with depth in Lake Fryxell and Lake Vanda.

Lake Hoare and both lobes of Lake Bonney were slightly enriched in Rb relative Cl in their surface waters. The Sr:Cl values in Lake Hoare and West Lake Bonney were nearly constant throughout their water columns. East Lake Bonney and Lake Fryxell had lower Sr:Cl values in their hypolimnions compared to their epilimnia, but water in Lake Vanda showed the opposite trend. Ba:Cl values remained constant throughout the profiles of Lake Fryxell and Lake Hoare.

In Lakes Bonney and Vanda, these ratios decreased with depth.

Don Juan Pond had the highest Li (2.65 x 105 μg/L), Sr (1.04 x 106 μg/L), and Ba (1.4 x 103

μg/L) concentrations of all the MCM lakes (Appendix D). The highest Rb concentration (294

μg/L) was found 35 m below the surface of E. Lake Bonney. Of all the MCM lakes, Lake Hoare, the freshest lake in this study, had the lowest Li (2.47 μg/L), Rb (2.11 μg/L), Sr (150 μg/L), and Ba

(1 μg/L) concentrations at 4 m. Water at and below the chemocline in Lakes Bonney and Vanda and the water in Don Juan Pond were enriched in Li (16-1,500 times higher) and Sr (2.3-130 times higher) relative to seawater. Water from 22 m and 35 m in E. Lake Bonney and 37 m in W. Lake

Bonney had higher Rb concentrations than seawater (1.3-2.5 times higher). Compared to seawater, barium concentrations were higher in water from 11-18 m in Lake Fryxell, Don Juan

Pond, and in and below the chemoclines in Lakes Bonney and Vanda (1.4-99 times higher). In all lakes, Li, Rb, and Sr increased in concentration with depth. Barium concentrations increased with depth in Lakes Hoare, Fryxell, and Vanda, but the hypolimnetic waters of both lobes of Lake

Bonney had concentrations lower than samples from the chemocline.

100 1000000

100000

10000

1000

SK g/L)

μ 100 GB Li ( Li 10 MCM 1 SW 0.1 y = 0.14x0.93 R² = 0.81 0.01 0.1 1 10 100 1000 10000 100000 1000000 Cl (mg/L)

Figure 3.4 Chloride and lithium concentrations in the lakes in this study. SK=Saskatchewan, GB=Great Basin, MCM=McMurdo Dry Valleys, SW=seawater (PILSON, 1998).

1000

100

SK g/L)

μ 10

GB Rb ( Rb

MCM 1 y = 0.40x0.49 SW R² = 0.81 0.1 0.1 1 10 100 1000 10000 100000 1000000 Cl (mg/L)

Figure 3.5 Chloride and rubidium concentrations in the lakes in this study. SK=Saskatchewan, GB=Great Basin, MCM=McMurdo Dry Valleys, SW=seawater (PILSON, 1998).

101 10000

1000

SK g/L)

μ 100 GB Ba ( Ba MCM

10 SW

1 10 100 1000 10000 100000 1000000 Cl (mg/L)

Figure 3.6 Chloride and barium concentrations in the lakes in this study. SK=Saskatchewan, GB=Great Basin, MCM=McMurdo Dry Valleys, SW=seawater (PILSON, 1998).

1000000

100000

SK 10000 GB

Sr (ug/L) Sr 1000 MCM

SW 100

10 1 10 100 1000 10000 100000 1000000 Cl (mg/L)

Figure 3.7 Chloride and strontium concentrations in the lakes in this study. SK=Saskatchewan, GB=Great Basin, MCM=McMurdo Dry Valleys, SW=seawater (PILSON, 1998).

102

10000000

1000000

100000 SK

g/L) GB μ 10000

Sr ( Sr MCM 1000 SW 100 y = 33.02x0.81 R² = 0.80 10 1 10 100 1000 10000 100000 1000000 Ca (mg/L)

Figure 3.8 Calcium and strontium concentrations in the lakes in this study. SK=Saskatchewan, GB=Great Basin, MCM=McMurdo Dry Valleys, SW=seawater (PILSON, 1998).

In the Saskatchewan lakes, Ceylon Lake had the highest concentrations of Li (1.31 x 104

μg/L), Rb (98.8 μg/L), and Sr ( 3.18 x 103 μg/L). Little Manitou Lake had the highest concentration of Ba (54 μg/L). Deadmoose Lake had the lowest concentrations of Li (1.9 x 103 μg/L), Rb (20.2

μg/L), and Ba (16 μg/L), and Freefight Lake had the lowest Sr concentration (35.1 μg/L).

Compared to seawater, the lakes in Saskatchewan had higher concentrations of Li (11-74 times higher) and Ba (1.1-3.9 times higher), and lower concentrations of Rb (1.2-6 times lower) and Sr

(2.6-232 times lower).

In the Great Basin, Mono Lake had the highest concentrations of Li (9.43 x 103 μg/L) and

Rb (1.63 x 103 μg/L) and Goose Lake has the lowest (5.90 μg/L and 8.60 μg/L, respectively). Goose

Lake also had the lowest Ba concentration (8 μg/L). Walker Lake had the highest concentration of Ba (160 μg/L) and Sr (4.85 x 103 μg/L). The Sr concentrations of Mono Lake and Abert Lake were below the detection limit of 0.05 µg/L.

103 3.5.2 PHREEQ

Although the model calculates the saturation states of a large suite of minerals, I sought to examine the only the simple binary salts of all the elements of interest which included, anhydrite (CaSO4), aragonite (CaCO3), barite (BaSO4), calcite (CaCO3), celestite(SrSO4), magnesite(MgCO3), strontianite (SrCO3) and witherite (BaCO3). The hydrated salts, gypsum

(CaSO4∙2H2O) and mirabilite (Na2SO4·10H2O) play a critical role in the evolution of many saline waters and are also of interest in this work. The activity products of these minerals are presented in Appendix B. I calculated saturation indices for the Great Basin and Saskatchewan lakes at two end member temperatures, 0°C and 20°C in order to examine how seasonal temperature variations might affect mineral precipitation within the lakes.

The Canadian Lakes wer all super saturated with respect to magnesite at 0°C and 20°C

(Appendix H). Waldsea Lake was supersaturated with aragonite, barite, and calcite at 0°C and

20°C. Little Manitou and Deadmoose Lakes were supersaturated with barite at 0°C and 20°C, supersaturated with respect to calcite at 20°C, and were in near equilibrium with aragonite at

20°C. At 0°C, Little Manitou and Deadmoose Lakes were undersaturated with calcite and aragonite. With respect to barite, Freefight and Ceylon Lakes were in near equilibrium at 20°C and supersaturated at 0°C. At 0°C, Ceylon Lake was the only lake in this study to be supersaturated with mirabilite.

All of the lakes sampled in the Great Basin were supersaturated with respect to calcite and magnesite at 0°C and 20°C (Calculations for Abert Lake were not attempted due to the lack of Ca data). All lakes except Pyramid Lake were supersaturated with aragonite at 0°C and 20°C.

Pyramid Lake was supersaturated with aragonite at 20°C, but it was in approximate equilibrium at 0°C. Pyramid Lake was also near equilibrium with respect to barite at 20°C. Walker and

104 Goose Lakes were supersaturated with strontianite at 0°C and 20°C, and strontianite was in near equilibrium at 20°C in Pyramid Lake.

The lakes in the McMurdo Dry Valleys have been sampled in such a way that their saturation indices in profile could be examined at in-situ temperatures (Figure 3.9, Figure 3.10).

Lake Hoare was in near equilibrium with barite at 30 m. Calcite was in near equilibrium in Lake

Fryxell at 8 m and from 12 to 14 m, and barite approached saturation at the sediment-water interface. Barite was in near equilibrium in the bottom water of Lake Vanda. In East Lake

Bonney, barite and magnesite were supersaturated at 22 and 35 m, and calcite and celestite were in near equilibrium at 35 m. In West Lake Bonney, barite was supersaturated at only 15 and 35 m.

3.6 Discussion

Normalizing solute concentration relative to chloride can be used to monitor the progression of evaporation and chemical evolution (EUGSTER and JONES, 1979). Chloride is considered conservative, that is, it is not removed from solution even in highly saline waters.

Alkali and alkaline earth elements have poor polarizability and tend to remain in solution since they are strongly attracted to water rather than inorganic ligands (STUMM and MORGAN, 1996).

Although these elements are highly mobile, their concentration in solution has been shown to be dependent on the age and type of rock (AIUPPA et al., 2000; GISLASON et al., 1996; GISLASON and

EUGSTER, 1987). Lithium and rubidium show strong positive relationships with chloride indicating that their behavior is conservative during brine evolution (Figure 3.4, Figure 3.5).

These results indicate that the concentration of lithium and rubidium in closed basin lakes is determined by the rock-water interactions in source waters followed by evapoconcentration and that no measurable removal has occurred at the salinities encountered in these lakes.

105 3.6.1 Inorganic mineral precipitation

Strontium shows a poor relationship with chloride (Figure 3.7), particularly in the sulfate and carbonate lakes indicating that strontium does not behave conservatively and is removed from solution during evolution. In saline lakes such as Pyramid Lake, especially during low-flow conditions, Ca, carbonate, Sr, and Ba can precipitate and be deposited into lake sediments (YANG et al., 2003). Strontianite is more soluble than calcite and is therefore predicted to precipitate during the early stages of brine formation. Strontium could form strontianite (SrCO3) or celestite

(SrSO4) during evapoconcentration and geochemical evolution. In addition, Sr can be incorporated into CaCO3, especially aragonite. Because the Saskatchewan lakes remain undersaturated with respect to strontianite 0° C and 20° C it is unlikely that strontium is being removed as either of these minerals (Figure 3.12). Likewise, data from the McMurdo Dry Valley lakes indicate that, the conditions for strontianite precipitation are thermodynamically unfavorable. Walker and Goose Lakes, however, are both supersaturated with respect to strontianite at 0° C and 20° C, and Pyramid Lake is in near equilibrium with strontianite at 0° C.

Given this information, it is possible that strontium is removed in Walker and Goose Lakes by strontianite precipitation throughout the year.

Celestite has been shown to be supersaturated in a hypersaline brines in the

Mediterranean Sea and is thought to be a primary strontium removal mechanism in that system

(DE LANGE et al., 1990b; KRUMGALZ et al., 1999). The geologic sedimentary rock record also provides insight into the strontium removal mechanisms. Celestite has been observed as druzes, fracture fillings, or cements, and crystals incorporated into the calcite matrix or encrusted on fossil fragments (SONNENFELD, 1984). In all of the lakes examined in this work, celestite is undersaturated with the exception of the bottom waters of East Lake Bonney where it is in near equilibrium (Figure 3.12). Depletion of barium and strontium is coupled with the calcium 106 depletion of a brine, and a continuous series of mineral precipitation from barite to celestite exists

(RUSHDI et al., 2000; SONNENFELD, 1984). In addition to being a removal “sink” for Ba, barite crystals in the ocean contain various amounts of Sr, and may exceed up to 20% (DEHAIRS et al.,

1980).

The poor relationship of barium to chloride indicates that barium is being removed in some of the lakes (Figure 3.6). Barite has a relatively low solubility and is supersaturated in surface waters such as the Mediterranean Sea and the brine-seawater interface of the Orca Basin in the Gulf of Mexico (DE LANGE et al., 1990b; KRUMGALZ et al., 1999; SCHIJF, 2007). In the open ocean, Ba distributions are primarily influenced by primary production and the dissolution and formation of barite in the water column leading to complex barium concentration profiles (CHAN et al., 1977; DEHAIRS et al., 1980; DEHAIRS et al., 1987; FALKNER et al., 1993). My thermodynamic modeling shows that barite is supersaturated in some lakes investigated here (Figure 3.12). In

Saskatchewan, Deadmoose, Little Manitou, and Waldsea Lakes are supersaturated at 20°C and

Ceylon and Freefight Lakes are in near equilibrium at this temperature. However, at near- freezing temperatures, all lakes become supersaturated with respect to barite suggesting that a seasonal precipitation of barite may be an important control on the barium concentration sulfate lakes. Winter removal of barite may also be occurring in Pyramid Lake in the Great Basin as it is in equilibrium at 20°C but is supersaturated at 0°C. The geochemical profiles of the McMurdo

Dry Valley Lakes give insight into barite removal processes in stratified chloride-type brines.

Lake Hoare and Lake Vanda are undersaturated with respect to barite, but the lake waters approach equilibrium with depth. The surface waters of both lobes of Lake Bonney are undersaturated in barite but become supersaturated at and below the chemocline. These results suggest that inorganic precipitation of barite may be an important removal mechanism in both the chloride and sulfate lakes with higher TDS. 107

Aragonite (CaCO3) Calcite (CaCO3)

-2 0 2 -2 0 2 0 0 10 10 20 20 30 30 40

40 (m) Depth

Depth (m) Depth 50 50

60 60

70 70

Strontianite (SrCO3) Witherite (BaCO3) -3 0 3 -6 0 6 0 0

10 10

20 20

30 30

40 40

Depth (m) Depth Depth (m) Depth

50 50

60 60

70 70 Lake Hoare Lake Fryxell W. Lake Bonney E. Lake Bonney Lake Vanda Don Juan Pond

Figure 3.9 Carbonate mineral saturation indices for the MCM lakes

108 Gypsum (CaSO4•H2O) Anhydrite (CaSO4) -4 0 4 -4 0 4 0 0

10 10

20 20

30 30

40 40

Depth (m) Depth Depth (m) Depth

50 50

60 60

70 70

Celestite (SrSO4) Barite (BaSO4) -4 0 4 -2 0 2 0 0

10 10

20 20

30 30

40 40

Depth (m) Depth Depth (m) Depth

50 50

60 60

70 70 Lake Hoare Lake Fryxell W. Lake Bonney E. Lake Bonney Lake Vanda Don Juan Pond

Figure 3.10 Sulfate mineral saturation indices for the MCM lakes.

109 Mirabilite (Na2SO4 ∙ 10H2O) -10 0 10 0

Lake Hoare Lake Fryxell W. Lake Bonney E. Lake Bonney Lake Vanda Don Juan Pond

Figure 3.11 Mirabilite saturation indices for the MCM lakes.

Although the solubility of witherite (BaCO3) is low, it is rarely observed as a primary mineral, but rather as an alteration product of barite (SONNENFELD, 1984). The calculations indicate that witherite is undersaturated in all lakes (Figure 3.12). A process such as this where carbonate is removed from solution and sulfate is solubilized might indirectly affect the saturation state of carbonates and sulfate minerals containing minor alkaline earth elements.

3.6.2 Replacement and incorporation into other minerals

As noted earlier, the precipitation of calcite is the first geochemical divide encountered during most brine evolution (Figure 3.1) (HARDIE and EUGSTER, 1970). The divalent elements strontium (Kd = 0.182) and barium (Kd = 0.0195), have low Kd values and therefore readily

110 substitute into calcite (RIMSTIDT et al., 1998). Furthermore, these minerals may be substituted for calcium in aragonite or gypsum, the former of which accommodates two substitute ions. It has also been shown that inorganic aragonite, particularly at lower salinities, can accommodate approximately 4.5 times more Li than calcite (MARRIOTT et al., 2004) After initial precipitation of calcite, a brine rich in calcium and/or magnesium, may evaporate enough to reach concentration where gypsum/anhydrite precipitation occurs. In very small amounts of strontium substitution occurs in gypsum and anhydrite, but this substitution is more pronounced during rapid precipitation (SONNENFELD, 1984).

With the exception of Abert Lake, which has no detectable alkalinity, the Great Basin lakes, are supersaturated with respect to calcite, and the saturation indices are higher at 20° C than at 0° C. Walker, Mono, and Goose Lakes are supersaturated with aragonite at 0° C and 20°

C, but Pyramid Lake only supersaturated at 20° C (Appendix B). In Mono Lake, calcite and aragonite precipitate around calcium enriched springs discharging into the carbonate rich lake water (JEHL, 1983). In addition to aragonite and calcite, metastable ikaite (CaCO3∙6H2O) precipitates along the shore during the winter months (BISCHOFF et al., 1993b). During the warmer months, ikaite is transformed into either CaCO3 or gaylussite (Na2Ca(CO3)2∙5H2O)

(BISCHOFF et al., 1993b). Sr and/or Ba may replace Ca in ikaite and be retained as it is transformed into calcite, aragonite, vaterite, or gaylussite. Strontium replacement in calcium carbonate minerals, especially aragonite, may be an important mechanism for strontium removal. However, incorporation into calcium sulfate is not a factor due to the undersaturation of gypsum and anhydrite in the Great Basin lakes.

Given the near freezing temperatures in the McMurdo Dry Valley lakes, ikaite precipitation may act as a mechanism for the removal of minor alkaline earth elements. Most occurrences of ikaite have been found in cold, anoxic marine sediments and debatably linked to 111 high orthophosphate concentrations (BISCHOFF et al., 1993a; OMELON et al., 2001). The water below 9 meters in Lake Fryxell is anoxic, and dissolved phosphorous concentrations increase with depth in the lake (GREEN et al., 1989). These qualities in addition to temperatures circa 2°C at the bottom of Lake Fryxell may provide favorable conditions for ikaite formation. The ice- covered dry valley lakes, Vanda, Bonney, Hoare, and Fryxell may present additional opportunities for ikaite precipitation as it has been shown to precipitate in hypersaline inclusions in Antarctic sea ice (DIECKMANN et al., 2008). Unfortunately, I was unable o include ikaite in the mineral saturation calculation.

Alternatively, some lakes in the MCM may have undergone freezing pathways as has been shown in the Vestfold Hills, Princess Elizabeth Land (STARK et al., 2003). Seawater freezing experiments has shown to ultimately produce antarcticite (CaCl2 · 6H2O) under the “Ringer-

Nelson-Thompson” parameters or MgCl2 · 12H2O under the “Glitterman pathway” (MARION et al., 1999). Although the Pitzer model in this work shows the MCM lakes to be undersaturated with respect to mirabilite (Figure 3.11), mirabilite and gypsum have shown to play a critical role in evolution of Cl brines under freezing conditions (MARION, 1997; MARION and FARREN, 1999;

MARION et al., 1999). Under freezing conditions, mirabilite (Na2SO4 · 10H2O) is the first mineral to precipitate at -7.3° C (MARION et al., 1999). Between -22.2° C and -22.9° C, mirabilite dissolves and gypsum (CaSO4 · 2H2O) and hydrohalite (NaCl · 2H2O) precipitate (MARION et al., 1999). The lowest temperature observed in the MCM lakes is -3.91° C in West Lake Bonney, indicating that the current temperatures in the MCM lakes are unfavorable for mirabilite precipitation.

However, the Pleistocene climate in the dry valleys climatic may have been ideal conditions for mirabilite formation as it has been demonstrated that the west lobe of Lake Bonney shows chemical similarities to Canadian and Fennoscandinavian Shield brines, both of which have

112 developed from seawater freezing (LYONS et al., 2005). Under these conditions, freezing seawater would become enriched in SO4, which is not the case for any of the MCM lakes (Appendix G).

The McMurdo Dry Valley Lakes are predominately undersaturated with respect to calcite and aragonite, thus direct incorporation of Ba or Sr into aragonite or calcite crystals is unlikely. They are also undersaturated in the sulfate minerals, gypsum and anhydrite, but in some instances, supersaturated with barite. This suggests that strontium and barium are not being removed by ionic substitution of calcium in calcite, aragonite, gypsum, or anhydrite.

The Canadian lakes in this study are mostly undersaturated with respect to calcite with the exception of Waldsea Lake (at 20° C and 0° C) and Little Manitou and Deadmoose Lakes at

20° C. This suggests that seasonal whitings of calcite could be a sink for minor elements.

Aragonite has been attributed to whiting events in Waldsea and Deadmoose Lakes when the lakes are at moderate salinity and, during higher salinity periods, the lakes are capable of producing gypsum, aragonite, magnesite, and dolomite precipitates (LAST and SCHWEYEN, 1985;

LAST and SLEZAK, 1986). According to the PHREEQ calculations, magnesite is supersaturated in all of the Saskatchewan lakes at high and low temperatures, but only Waldesa Lake is supersaturated with respect to aragonite (Appendix H). The undersaturation of gypsum, dolomite, and aragonite in some lakes is likely due to the seasonality of the salinity, temperature, and chemical composition of the lakes in southern Saskatchewan. These data indicate that the removal of minor elements through the precipitation of gypsum, aragonite, and dolomite may vary spatially and temporally.

113

A) 3 B) 6 II I II I

Calcite 0 0 -1.5 0 1.5 -1.5 0 1.5

III IV III IV -3 -6 Strontianite Witherite

Canada 20C Canada 0C Great Basin 20C Great Basin 0C MCM C) 4 D) 4 II I II I

Gypsum 0 0 -7 0 7 -7 0 7

III IV III IV -4 -4 Celestite Barite

Figure 3.12 The saturation indices of selected minerals in the lakes of this study. A) Quadrant I: calcite and strontianite are supersaturated. Quadrant II: strontianite is supersaturated, and calcite is undersaturated. Quadrant III: strontianite and calcite are undersaturated. Quadrant IV: calcite is supersaturated, and strontianite is undersaturated. B) Quadrant I: calcite and witherite are supersaturated. Quadrant II: witherite is supersaturated, and calcite is undersaturated. Quadrant III: witherite and calcite are undersaturated. Quadrant IV: calcite is supersaturated, and witherite is undersaturated. C) Quadrant I: gypsum and celestite are supersaturated.

114

Because of their similar charges, ionic radii, and electro negativities, Rb may replace K in crystal lattices. As such, carnallite and sylvite may incorporate Rb, but the precipitation of these ions is limited due to their very high solubility. For example, during the final stages of seawater evaporation, Rb preferentially concentrates in carnallite at concentrations of up to 300 ppm

(HOSLER, 1979). All lakes in this study are grossly undersaturated with respect to carnallite and sylvite, thus Rb is not being removed through the precipitation of these minerals.

All lithium salts are extremely soluble and as a result, lithium does not readily precipitate until the end stages of brine evolution. During these final stages, the highly soluble salts such as the chlorides reach saturation. The solubility of lithium is so high that it concentrates beyond bischofite precipitation and may be incorporated in halite, polyhalite, carnallite, sylvites, and borates (Lepeshkov et al., 1970; Kropachev et al., 1972; Zherebtsova and Volkova, 1966). At some point during the end stages and halite precipitation, all strontium is removed from the brine as

SrCl because the solubility of SrCl decreases dramatically in the presence of NaCl (SONNENFELD,

1984). Although the lakes in this study are at various stages of brine evolution, there is no evidence to suggest that Rb or Li are being removed from the lakes as binary salts or that SrCl is precipitating.

3.6.3 The potential importance of biogenic precipitation

Celestite precipitation may occur inorganically or through biological activity. In freshwater systems, cyanobacteria has been shown to mediate the formation of celestite and strontianite (SCHULTZELAM and BEVERIDGE, 1994). Strontium depletion in the upper ocean has been ascribed to the production of a celestite skeleton by surface-dwelling acantharia (BERNSTEIN et al., 1992). The death and decomposition of these organisms can also facilitate in the precipitation of barite (BERNSTEIN and BYRNE, 2004). A strong correlation between barium and

115 nutrient concentration in the surface water of the oceans suggests biotic control of Ba distribution in seawater (DEHAIRS et al., 1980; MOORE and FALKNER, 1999). Although barite production is common in the open ocean, it functions as more of a transport mechanism rather than a permanent sink of barium as approximately 85% is released into the deep ocean during the breakdown and mineralization of organic matter (WOLGEMUTH and BROECKER, 1970). This dissolution of barite is also evidenced by a small barium concentration maximum at the O2/H2S interface in the Black Sea (FALKNER et al., 1993). Strontium also has been shown to be influenced by biogeochemical processes in the marine environment and in some lacustrine environments.

Although these ecosystems are quite different than the lakes in this study, biologically mediated transport cannot be ruled out as a process occurring in the systems studied here.

3.7 Conclusions

This study has elucidated potential controls on the high concentrations of minor elements in closed-basin lakes. In all the anionic water types, lithium and rubidium appear to behave conservatively and are the result of solute input and evapoconcentration. Unlike the conservative behavior of lithium and rubidium, strontium and barium do not behave conservatively, and thermodynamic calculations suggest they are removed by different mechanisms depending on brine type. This study has expanded upon the Eugster-Hardie model by incorporating minor elements into geochemical pathways (Figure 3.1). I have shown that intermediate steps located within path I may include strontianite precipitation and barite may form along path II. Strontium appears to be depleted in sulfate and carbonate brines. In the carbonate brines of the Great Basin, strontium may be removed as strontianite (SrCO3) or substituted for calcite in calcite, ikaite, or especially aragonite. Strontium may be removed from the sulfate-type lakes of Saskatchewan by substitution into calcium carbonates as these lakes are

116 supersaturated with respect to calcite and aragonite. Gypsum is a phase that is not important for

Sr in these lacustrine systems as it is undersaturated in all of the lakes under investigation. Barite is supersaturated in the McMurdo Dry Valley lakes (chloride-type) and the Saskatchewan lakes indicating that it a possible important sink in these systems. Barite and celestite have been shown to be biologically precipitated, and the role of biogenic removal of barium and strontium in these saline lakes warrants further investigation.

3.8 Acknowledgements

I wish to thank J. Olesik and A. Lutton of the Trace Element Research Laboratory at The

Ohio State University for instrumentation usage. J.C. Priscu and his MCM-LTER limnology team collected MCM lake samples. W. Last generously contributed samples from Saskatchewan to this work. S. Fortner and J. Barker assisted in GB and MCM sample collection. K.A. Welch, C.

Gardner, and C. Maxwell assisted in major ion and alkalinity analysis. I am also indebted to the following people for their thoughtful comments: W. Berry Lyons, Ozeas Costa, Bill Green, Anne

Carey, and Dave Porinchu. The support of many from the MCM-LTER program, Raytheon, PHI, and DRI made this work possible. This work was funded by NSF-ANT grant #0423595

117

Chapter 4 Conclusions

The overall objective of this dissertation was to evaluate the behavior of alkaline earth elements and alkali metals in endorheic lakes. This work examined the concentrations and relationships of minor alkali elements in inflow streams and terminal lakes of the McMurdo Dry

Valleys, the lithium isotopic dynamics in the aqueous systems of the dry valleys, and the potential removal mechanisms of these elements in three brine types. This work has lead to the following conclusions:

Hypothesis 1a: Like the major ions in MCM streams, the sources of minor alkali metals and

alkaline earth elements are the result of precipitation, chemical weathering, and salt dissolution

 Lithium has a strong correlation with chloride in the streams indicating that

lithium is not being removed from the streams.

 Strontium and calcium are being weathered at the same rate and may be coming

from the same source.

 Rubidium and barium have complex behaviors in streams, perhaps being

removed through adsorption onto clay minerals.

 The dissolution of marine salts is not the sole source of Li, Rb, Sr, and Ba to the

lakes. Chemical weathering likely contributes a significant amount of minor

cations to the streams and lakes.

118 Hypothesis 1b: Biological uptake and adsorption onto clays are sinks for minor alkalis in MCM streams.

 There is not enough evidence to show that biological uptake is a significant sink

of alkalis, however, longitudinal profiles suggests an unknown mechanism for

removing barium.

 Lithium and strontium appear to be conservative in streams, and rubidium is

likely removed due to clay adsorption.

Hypothesis 2a: Lithium isotopic ratios will identify the source of lithium to the MCM streams.

 The isotopic mixing curve generated by the glacier, stream, and lake data

indicate that primary precipitation and the dissolution of marine salts are the

dominant contributors of Li the the hydrologic system in the dry valleys..

Hypothesis 2b: The fresher surface waters of the lakes will be isotopically similar to the streams

reflecting the sources on the modern lithium inputs.

 Lakes Hoare and the surface water of Lake Fryxell represent end members of the

precipitation-seawater mixing line. Thus, the epilimnia of these lakes are similar

to stream water that has acquired more solutes through salt dissolution and

chemical weathering.

Hypothesis 2c: The older hypersaline waters will have unique signatures as a result of the climatic

history of the lake..

 The hypolimnion of Lake Fryxell has a strong seawater signature reaffirming the

understanding that the hypolimnion is of marine origin.

119  The monimolimnetic waters of Lake Bonney are derived from past dilution of

seawater by glacial melt.

 The solutes in the surface waters of Lake Vanda are the result of Onyx River

water and diffusion from saline bottom waters.

 Don Juan Pond is not characterized by evapoconcentration of glacial melt, rather

its water source is most likely groundwater.

Hypothesis 3a: Mineral saturation indices will reflect the degree of brine evolution

 In all the anionic water types, lithium and rubidium appear to behave

conservatively and are the result of solute input and evapoconcentration. These

ions do not play a straightforward role in mineral saturation.

 The more saline lakes in the McMurdo Dry Valleys have higher saturation

indices than the more dilute lakes in the dry valleys.

 Lakes that are highly evolved may be completely depleted in a minor ion. Such

is the case for Mono Lake and Abert Lake which have no detectable strontium.

Hypothesis 3b: Brines of different major anion type will display differences in major and minor mineral saturation indices

 Strontianite, calcite, and aragonite are supersaturated in some carbonate-type

lakes.

 Barite is supersaturated in most sulfate and chloride lakes and may be the only

sulfate mineral precipitating in the Saskatchewan Lakes.

Hypotheses 3c: Examining the saturation indices of major and minor carbonate and sulfate minerals will indicate potential sinks for alkaline earth elements and alkali metals in closed systems. 120  Strontium appears to be depleted in sulfate and carbonate brines. In the

carbonate brines of the Great Basin strontium may be removed as strontianite

(SrCO3) or substituted for calcite in calcite, ikaite, or especially aragonite.

Strontium may be removed from the sulfate-type lakes of Saskatchewan by

substitution into calcium carbonates as these lakes are supersaturated with

respect to calcite and aragonite.

 Barite is supersaturated in some McMurdo Dry Valley lakes (chloride-type) and

the Saskatchewan (sulfate-type) lakes indicating that it a possible sink in these

systems.

This dissertation has shed important light onto the behavior of minor elements in selected lakes in Antarctica and North America. This should be considered a preliminary study as there is much work to be done on the behavior of these elements in various settings. I suggest that future work may consider: examining the chemical composition of detrital and authigenic sediments, investigating uptake of barium in Antarctic streams, determining the seasonality of minor elemental removal from lakes, conducting research on various brine types including mixed-types.

121

Appendix A: Map of the McMurdo Dry Valleys

122 Wright Valley

Taylor Valley

123

Commonwealth Str.

Lost Seal Str

Huey Cr. McKnight Cr. Andrews Cr. Aiken Cr. Canada Str. Andersen Cr. Von Guerard Str. Harnish Cr. Bowles Cr. House Str. Green. Cr.

1 McKay Cr. Mariah Cr. 2 Crescent Str. 4 Wharton Cr. Delta Str.

Priscu Str.

Lawson Cr. Santa Fe Lyons Cr.

124

Appendix B: Detection Limits, Precision Data, and Blank Concentrations

125 DL (μg/L) Na K Mg Ca F Cl SO4 Li Rb Sr Ba Cs

Glaciers 0.01 n.d.

Streams 0.03 0.07 0.09 0.01 0.05

Lakes 0.01 0.01 0.05 0.003 0.05

Detection limits of stream and lake samples.

Na K Mg Ca F Cl SO4 Li Rb Sr Ba Cs

Maximum variability of duplicate lake and stream samples (%)

Lakes 0.6 7.5 3.7 0.6 2.5 3.1 5.1 6 7 8 14

Streams 0.5 0.8 0.7 0.4 5.6 0.5 1.0 8 16 5 5

Maximum variability of sample replicates (%)

All Locations 0.6 3.7 0.6 0.6 2.5 3.1 5.1

Relative standard deviation of check standards (%)

Lakes 1.8 0.7 0.7 0.7 7.7

Streams 4.2 0.6 0.6 1.9 7.7

Maximum variability of duplicate lake and stream samples.

Lake Depth (m) Difference (%)

Hoare 6-30 1-11

Fryxell 10-18 4-10

E. Bonney 8-37 1-10

W. Bonney 6-37 1-19

Relative differences in alkalinity between replicated of MCM samples collected below chemoclines.

126

Na K Mg Ca F Cl SO4 Li Rb Sr Ba Cs Stream Blank DL DL DL 0.05 DL 0.03 DL DL DL 0.04 0.34 DL 1 Stream Blank DL DL DL 0.02 DL DL DL 0.06 DL DL DL DL 2 Lake Hoare 0.1 0.1 0.1 0.1 DL DL DL DL DL DL 0.1 DL

27 Lake Fryxell 2.3 0.2 0.4 0.5 DL 0.9 DL 0.6 DL 0.2 0.1 DL

East Bonney 3.9 0.2 2.5 0.3 DL 5.8 1.2 DL DL 0.3 0.7 DL

West Bonney 0.8 DL 0.2 0.2 DL 1.5 0.5 DL DL 0.2 0.4 DL

Lake Vanda 0.1 DL 0.1 0.1 DL 1.3 0.1 0.3 DL 0.2 DL DL

Great Basin DL DL DL DL DL DL DL 0.5 DL DL DL n.d.

Major and minor ion concentration in field blanks. Major ions, Na, K, Mg, Ca, F, Cl, and SO4, are in mg/L. Minor ions, Li, Rb, Sr, and Ba, are in μg/L. DL = detection limit. n.d. = not determined.

127

Appendix C: Mean Major and Minor Ion Concentrations of Streams

128

Cl SO4 Na K Mg Ca Li Rb Sr Ba Alk pH

mg/L mg/L mg/L mg/L mg/L mg/L ug/L ug/L ug/L ug/L meq/L

Aiken 20.7 6.32 30.8 5.15 4.04 13.5 2.84 0.71 60.4 0.2 1.75 n.d. 17.9 7.41 25.9 3.69 2.79 9.35 3.43 0.92 80.8 0.2 1.25 n.d.

17.8 6.00 23.4 3.23 3.03 9.73 2.45 0.79 60.6 0.7 1.52 n.d.

21.8 6.61 28.8 3.76 3.09 9.93 2.88 1.00 62.8 3.4 1.49 n.d.

16.4 9.30 17.4 3.22 3.33 12.8 2.16 0.78 66.2 0.5 1.20 n.d.

22.5 11.1 23.0 4.15 4.17 15.7 2.46 0.89 81.1 0.6 1.48 n.d.

20.1 7.38 23.5 3.46 3.72 11.5 2.57 0.74 64.8 1.5 1.34 n.d.

Mean 19.6 7.73 24.7 3.81 3.45 11.8 2.68 0.83 68.1 1.0 1.43 n.d. Andersen 2.23 2.15 1.21 0.42 0.37 2.36 0.39 0.38 15.8 0.5 0.10 n.d. 1.79 1.92 0.96 0.37 0.37 2.12 0.28 0.32 13.9 0.5 0.10 n.d.

1.94 2.04 1.04 0.41 0.40 2.67 0.36 0.37 17.0 1.3 0.12 n.d.

29 2.20 2.50 1.23 0.48 0.44 3.48 0.43 0.42 21.3 1.0 0.16 n.d.

1.68 1.92 1.01 0.40 0.24 2.63 0.38 0.39 16.8 0.3 0.12 n.d.

1.41 1.61 0.81 0.33 0.19 2.15 0.31 0.33 13.9 0.6 0.09 n.d.

2.00 3.91 1.36 0.53 0.47 3.32 0.42 0.41 18.5 0.4 0.14 n.d.

1.93 3.72 1.28 0.52 0.45 3.14 0.39 0.43 17.8 0.4 0.13 n.d.

1.62 3.10 1.09 0.45 0.43 2.74 0.36 0.32 15.1 0.8 0.12 n.d.

4.68 6.49 3.05 0.93 0.69 6.90 0.90 0.67 42.3 0.9 0.29 6.35

6.18 8.77 3.76 1.12 0.91 8.64 1.02 0.79 52.2 3.6 0.33 6.46

3.07 4.13 1.98 0.63 0.44 4.43 0.61 0.50 27.9 2.2 0.18 6.52

2.75 3.75 1.78 0.51 0.38 3.75 0.52 0.41 23.8 0.6 0.15 7.60

2.31 3.10 1.34 0.39 0.30 2.96 0.28 0.24 13.0 0.4 0.11 6.67

2.12 2.16 1.09 0.30 0.24 2.08 0.30 0.34 14.8 0.4 0.07 6.89

1.93 1.96 0.97 0.28 0.23 2.20 0.26 0.34 14.9 0.4 0.08 6.96

129 2.50 2.60 1.27 0.40 0.33 3.44 0.38 0.42 22.0 0.6 0.14 6.98

3.80 3.84 1.92 0.64 0.56 6.14 0.59 0.61 37.9 0.9 0.26 7.11

2.66 2.86 1.54 0.49 0.35 3.82 0.52 0.50 24.2 0.7 0.16 7.06

2.39 2.72 1.53 0.49 0.34 3.68 0.53 0.51 22.8 0.4 0.16 7.04

3.01 3.89 2.06 0.78 0.59 6.59 0.78 0.71 39.3 0.7 0.32 7.02

2.35 3.06 1.53 0.54 0.39 4.36 0.56 0.53 26.7 0.4 0.20 6.91

1.82 2.28 1.18 0.39 0.28 2.99 0.42 0.41 18.9 0.3 0.13 6.80

3.36 4.59 2.59 0.85 0.64 7.29 0.82 0.75 44.3 0.7 0.36 7.04

2.72 3.69 1.89 0.65 0.50 5.47 0.64 0.51 34.3 0.5 0.26 7.04

2.21 2.82 1.44 0.47 0.33 3.54 0.49 0.45 22.6 1.3 0.15 6.81

1.86 2.55 1.17 0.37 0.27 2.85 0.42 0.38 18.3 0.4 0.12 6.67

1.66 2.47 1.02 0.36 0.26 3.03 0.41 0.38 19.1 0.4 0.13 6.77

1.61 2.37 0.99 0.36 0.26 2.97 0.37 0.39 19.2 1.2 0.13 6.70

1 1.76 2.54 1.04 0.38 0.28 3.18 0.41 0.39 20.0 0.4 0.13 6.85

0 1.92 2.88 1.14 0.41 0.31 3.69 0.46 0.42 23.4 0.4 0.15 6.82

2.18 2.52 1.33 0.52 0.39 4.72 0.54 0.50 29.3 0.6 0.22 6.89

2.44 1.62 1.49 0.60 0.46 5.53 0.63 0.57 33.9 1.2 0.29 6.90

1.99 2.99 1.20 0.45 0.34 4.09 0.48 0.45 25.6 0.4 0.18 6.80 2.01 2.69 1.38 0.54 0.46 3.69 0.47 0.42 22.1 0.4 0.18 n.d.

2.35 3.01 1.67 0.67 0.52 4.71 0.54 0.51 26.8 0.5 0.24 n.d.

2.16 3.29 1.46 0.56 0.49 4.13 0.48 0.43 23.7 0.9 0.19 n.d.

2.50 4.93 1.70 0.43 0.43 3.79 0.51 0.51 25.0 1.3 0.19 n.d.

3.38 6.00 2.09 0.56 0.52 4.32 0.49 0.55 24.0 2.2 0.26 n.d.

18.3 41.6 9.68 3.03 3.11 22.9 1.99 1.23 121 4.3 0.63 n.d.

3.97 6.70 2.56 0.62 0.61 5.60 0.75 0.62 35.4 1.5 0.25 n.d.

11.5 27.9 9.05 2.28 2.42 19.3 1.94 1.68 113 5.3 0.75 n.d.

130 3.53 6.74 3.26 1.12 0.89 8.99 1.06 0.85 51.5 1.2 0.50 n.d.

Mean 3.07 4.85 1.93 0.63 0.53 4.89 0.58 0.52 29.4 1.0 0.43 6.87 Andrews 1.24 0.78 2.09 0.66 0.41 3.93 0.95 0.86 54.4 5.5 0.28 7.27 Bowles 2.32 1.60 2.64 2.43 0.85 8.00 0.77 0.74 48.7 2.9 0.57 n.d. 0.80 0.81 1.27 1.03 0.41 5.01 0.48 0.65 27.6 1.4 0.40 n.d.

Mean 1.56 1.20 1.95 1.73 0.63 6.51 0.62 0.70 38.1 2.2 0.49 n.d. Canada 1.53 1.80 4.35 0.43 0.29 2.33 0.35 2.00 16.1 1.1 0.17 n.d. 1.76 2.11 1.09 0.44 0.30 2.20 0.31 0.72 15.2 1.6 0.14 n.d.

1.45 1.53 1.10 0.34 0.27 2.21 0.26 0.60 13.9 0.7 0.17 n.d.

1.13 1.39 0.86 0.43 0.26 2.08 0.33 1.32 15.2 1.8 0.16 n.d.

1.98 1.54 1.35 0.63 0.46 3.44 0.29 0.79 21.8 0.5 0.20 7.11

1.88 1.46 1.35 0.58 0.46 3.30 0.23 0.75 20.8 3.3 0.19 7.07

1.87 1.45 1.35 0.55 0.44 3.24 0.26 0.74 20.1 0.3 0.19 7.08

1 1.79 1.41 1.22 0.53 0.41 2.71 0.21 0.68 17.2 1.7 0.15 7.04

0.66 0.39 0.49 0.16 0.20 0.99 0.17 0.30 7.0 1.6 0.06 6.52

Mean 1.56 1.45 1.46 0.45 0.34 2.50 0.27 0.88 16.7 1.4 0.16 6.96 Commonwealth 4.33 2.33 3.04 0.52 0.57 2.20 0.56 0.49 14.4 3.6 0.21 n.d. 5.93 2.74 3.97 0.80 0.71 3.21 0.46 0.69 18.1 2.4 0.23 n.d.

5.11 2.56 3.76 0.53 0.84 3.52 0.50 0.34 18.0 0.3 0.28 n.d.

Mean 5.12 2.54 3.59 0.62 0.71 2.98 0.50 0.51 16.8 2.1 0.24 n.d. Crescent 9.38 5.51 6.72 2.04 2.33 25.8 1.59 1.41 153 1.1 1.52 n.d. 11.5 5.59 8.32 2.92 2.48 26.6 2.11 1.03 167 1.5 1.64 n.d.

16.6 8.98 9.07 3.40 3.12 32.2 2.32 1.25 183 2.8 1.71 n.d.

12.2 6.59 7.86 2.71 2.61 27.7 1.94 0.84 160 0.4 1.64 n.d.

14.6 6.79 10.2 3.11 2.66 25.3 2.18 2.73 164 0.7 1.56 n.d.

Mean 12.8 6.69 8.43 2.84 2.64 27.5 2.03 1.45 166 1.3 0.83 n.d.

131 Delta 26.8 9.31 11.7 2.74 3.58 21.8 2.43 1.03 163 1.3 1.01 7.94 27.3 9.94 10.1 2.53 3.45 25.8 2.03 1.45 190 2.5 1.22 n.d.

13.6 7.87 8.16 2.02 2.54 18.7 1.49 0.97 128 0.7 1.05 n.d.

14.3 7.35 7.68 1.84 2.46 18.6 1.74 1.06 138 2.9 1.07 n.d.

17.3 8.25 8.87 2.30 3.10 22.9 1.64 1.11 150 1.2 1.25 n.d.

21.2 9.63 9.92 2.64 3.34 24.8 2.24 2.58 183 2.7 1.32 n.d.

Mean 20.1 8.73 9.40 2.34 3.08 22.1 1.93 1.37 159 1.9 1.15 7.94 Green 2.42 1.17 2.10 0.84 0.70 5.65 0.85 0.53 49.5 0.3 0.36 7.29 1.96 0.81 1.58 0.73 0.45 3.97 0.78 0.62 34.0 1.1 0.31 n.d.

1.41 0.90 1.17 0.43 0.30 2.80 0.25 0.36 20.0 2.9 0.21 n.d.

1.44 0.82 1.31 0.55 0.38 3.61 0.22 0.34 26.5 0.6 0.28 n.d.

1.55 0.74 1.36 0.54 0.34 3.51 0.19 0.34 29.5 0.8 0.27 n.d.

Mean 1.75 0.89 1.50 0.62 0.43 3.91 0.45 0.44 31.8 1.2 0.69 7.29

1 32 Harnish 17.6 9.31 14.1 3.40 4.21 27.0 2.40 1.15 161 1.2 1.80 n.d.

22.4 12.9 19.5 5.25 4.63 28.8 3.37 1.35 183 2.1 2.22 n.d. 19.9 10.4 18.3 5.22 3.91 23.6 3.42 1.56 150 1.7 1.68 n.d.

Mean 20.0 10.9 17.3 4.63 4.25 26.5 3.06 1.36 165 1.7 1.90 n.d. House 2.94 4.42 1.60 0.75 0.68 13.1 0.84 1.91 79.7 1.5 0.62 7.94 2.93 3.36 1.45 0.38 0.50 7.73 0.52 1.10 43.0 1.5 0.46 n.d.

4.90 5.23 2.84 0.75 1.12 14.0 0.91 1.60 81.3 2.6 0.85 n.d.

3.79 4.70 1.95 0.60 0.69 10.4 0.86 1.49 62.6 2.5 0.62 n.d.

2.84 3.46 1.65 0.45 0.59 9.97 0.61 1.10 55.8 1.2 0.59 n.d.

Mean 3.48 4.23 1.90 0.59 0.72 11.0 0.75 1.44 64.5 1.9 0.63 7.94 Huey 37.0 36.2 12.8 3.45 4.05 34.0 5.01 1.66 182 6.0 1.07 n.d. 6.00 7.49 5.56 1.89 1.23 10.2 1.86 0.96 53.9 2.1 0.64 n.d.

15.7 19.8 10.3 3.48 2.98 24.3 3.50 1.50 126 2.8 1.22 n.d.

132 12.0 13.2 8.28 2.44 2.13 16.7 3.07 1.80 96.7 5.4 0.90 n.d.

Mean 17.7 19.2 9.23 2.82 2.59 21.3 3.36 1.48 115 4.1 0.96 n.d. Lawson 4.48 2.97 3.32 0.37 0.69 1.87 1.13 1.73 12.9 11 0.13 n.d. 3.56 3.14 2.33 0.23 0.45 1.50 0.24 0.35 8.0 0.3 0.11 n.d.

3.44 2.21 2.67 0.33 0.48 1.50 0.12 DL 5.7 1.5 0.14 n.d.

4.00 3.06 2.76 0.33 0.56 1.53 0.40 0.24 6.2 0.2 0.07 7.05

Mean 3.87 2.84 2.77 0.32 0.54 1.60 0.47 0.58 8.2 3.3 0.11 7.05 Lost Seal 6.00 2.30 6.22 1.57 1.25 6.86 1.41 0.67 37.4 0.4 0.73 n.d. 4.66 2.63 4.15 1.00 0.80 4.01 0.88 0.40 22.1 0.1 0.40 n.d.

28.2 12.8 10.4 1.91 3.45 15.5 1.57 0.81 102 3.7 0.73 n.d.

8.03 4.10 5.11 1.07 1.29 6.48 0.59 0.43 35.1 0.3 0.40 n.d.

11.1 8.52 8.28 1.60 2.03 10.4 1.31 0.61 57.5 0.4 0.66 n.d.

8.20 3.82 7.00 1.47 1.37 6.77 0.94 0.54 34.0 1.4 0.51 n.d. 1

3 Mean 11.0 5.69 6.86 1.44 1.70 8.34 1.12 0.58 48.0 1.0 0.52 n.d.

Lyons 3.67 13.3 2.58 0.29 0.49 7.96 0.43 DL 37.3 2.1 0.24 n.d. Mariah 1.46 1.03 0.97 0.28 0.20 1.63 0.09 0.24 12.4 0.2 0.34 n.d. McKay 0.61 0.90 0.46 0.13 0.13 4.95 0.15 0.70 34.1 4.8 0.30 n.d. 1.58 2.03 0.96 0.34 0.25 7.40 0.46 1.03 55.1 1.8 0.42 n.d.

1.91 3.29 0.90 0.32 0.28 9.04 0.43 1.02 65.0 2.3 0.45 n.d.

0.99 1.59 0.58 0.29 0.22 8.57 0.37 0.93 57.9 1.2 0.50 n.d.

1.16 1.58 0.64 0.33 0.26 8.06 0.38 1.07 58.2 1.5 0.39 7.36

0.84 1.31 3.51 0.18 0.17 6.70 0.28 0.68 46.5 1.5 0.38 n.d.

Mean 1.18 1.78 1.18 0.26 0.22 7.45 0.34 0.90 52.8 2.2 0.56 7.36 McKnight 10.1 3.49 15.9 1.60 4.63 15.6 1.58 0.56 46.5 0.9 1.00 n.d. 7.64 2.62 11.7 1.91 1.76 8.04 1.57 0.39 43.5 0.2 0.82 n.d.

6.96 3.16 11.6 1.74 1.67 6.89 1.39 0.44 34.1 1.3 0.90 n.d.

133 8.95 3.22 14.6 2.20 1.67 7.15 1.58 0.46 35.8 0.3 1.30 n.d.

Mean 8.40 3.12 13.4 1.86 2.43 9.42 1.53 0.46 40.0 0.7 1.01 n.d. Onyx 4.19 2.22 2.42 0.51 0.56 2.08 0.29 0.51 12.2 0.3 0.10 6.64 4.40 2.35 2.48 0.53 0.57 2.17 0.30 0.52 12.7 0.3 0.10 6.93

4.26 2.30 2.51 0.66 0.61 2.68 0.49 0.65 15.0 2.5 0.14 6.96

4.79 2.63 2.82 0.77 0.70 3.22 0.60 0.73 17.9 0.5 0.17 7.00

4.85 2.62 2.91 0.82 0.73 3.43 0.68 0.78 18.5 1.6 0.18 7.06

4.94 2.69 2.96 0.86 0.73 3.49 0.74 0.85 19.2 0.5 0.19 n.d.

5.46 3.08 3.49 1.04 0.88 4.06 0.89 0.66 21.1 0.8 0.23 7.16

8.02 4.82 5.02 1.24 1.27 5.62 1.17 0.65 28.6 0.6 0.30 7.26

Mean 4.57 2.47 2.68 0.69 0.65 2.84 0.52 0.67 15.9 1.0 0.18 7.00 Priscu 10.6 6.31 6.13 1.71 2.51 11.1 2.22 2.08 113 6.9 0.90 n.d. 10.5 7.36 6.17 1.70 2.70 9.96 1.74 1.48 83.0 1.5 0.66 n.d. 1

4 15.3 12.1 7.61 2.02 3.26 14.0 2.00 1.25 134 1.7 0.79 n.d.

30.6 12.3 9.76 2.68 5.93 19.4 2.87 1.66 160 2.0 0.82 7.71

Mean 16.7 9.52 7.42 2.03 3.60 13.6 2.21 1.62 123 3.0 0.79 7.71 Santa Fe 18.5 15.2 10.6 1.09 2.17 13.5 1.38 0.53 81.9 1.6 0.52 n.d. 12.4 20.7 15.8 1.57 4.73 15.7 4.72 4.53 128 16 0.77 n.d.

16.6 33.9 11.1 1.12 1.85 18.4 1.46 0.35 79.4 1.6 0.78 n.d.

13.4 45.0 10.0 1.23 1.96 20.7 1.40 0.37 93.9 4.9 0.68 n.d.

21.2 54.1 16.8 1.74 2.66 25.5 2.36 0.32 112 1.2 0.54 8.54

Mean 16.4 33.8 12.8 1.35 2.67 18.7 2.26 1.22 99.1 5.0 0.66 8.54 Von Guerard 8.12 5.00 8.82 3.04 2.20 18.7 2.04 1.46 104 2.1 1.27 n.d. 6.30 4.57 7.14 2.38 1.73 15.7 1.65 0.86 73.0 2.7 1.13 n.d.

7.75 5.19 8.57 2.67 2.20 19.3 1.78 1.11 98.4 0.7 1.32 n.d.

Mean 7.39 4.92 8.18 2.69 2.05 17.9 1.82 1.14 91.7 1.8 1.24 n.d.

134

Wharton 1.42 1.26 1.29 0.39 0.33 7.21 0.35 0.77 51.9 0.9 0.47 n.d. 2.20 2.53 1.43 0.60 0.41 10.1 0.75 1.22 78.2 1.6 0.57 n.d.

2.14 2.84 1.09 0.44 0.34 8.98 0.58 0.91 68.1 2.9 0.63 n.d.

5 1.42 1.83 0.90 0.44 0.39 10.3 0.66 0.80 73.8 1.2 0.60 n.d.

1.45 1.38 0.86 0.32 0.31 7.44 0.46 0.67 56.2 1.6 0.68 n.d.)

2.14 2.05 1.39 0.56 0.47 9.94 0.65 0.95 76.2 1.2 0.51 7.47

Mean 1.80 1.98 1.16 0.46 0.38 8.98 0.57 0.89 67.4 1.6 0.58 7.47

135

Appendix D: Minor Ion Concentrations for the Lakes in This Study

136

Li Rb Sr Ba Saskatchewan Lakes µg/l µg/l µg/l µg/l

Ceylon Lake 13,100 98.8 3,180 40

Deadmoose Lake 1,900 20.2 490 16

Freefight Lake 5,700 70.3 35.1 20

Little Manitou Lake 11,400 37.4 1,890 54

Waldsea Lake 2,900 25.1 1,800 38

Li Rb Sr Ba Great Basin Lakes µg/l µg/l µg/l µg/l

Mono Lake 9,430 1630 DL 40

Pyramid Lake 465 34.4 126 40

Walker Lake 5m 1,390 31.7 4,850 160

Abert Lake 67.2 261 DL 38

Goose Lake 5.90 8.60 132 8

137 Li Rb Sr Ba McMurdo Dry Valleys Lakes µg/l µg/l µg/l µg/l

Lake Hoare 4m 2.47 2.11 150 2

Lake Hoare 5m 4.06 2.16 228 3

Lake Hoare 6m 7.06 3.31 389 5

Lake Hoare 8m 9.75 4.37 549 7

Lake Hoare 10m 12.3 5.24 669 9

Lake Hoare 12m 13.2 5.55 714 10 Lake Hoare 14m

1 14.1 5.67 774 10

8 Lake Hoare 16m 15.9 5.96 865 12

Lake Hoare 18m 16.6 5.96 878 12

Lake Hoare 20m 16.1 5.90 883 12

Lake Hoare 22m 16.3 5.97 893 12

Lake Hoare 25m 16.5 6.01 904 12

Lake Hoare 30m 16.8 5.98 928 13

Lake Fryxell 6m 23.6 6.38 668 5

Lake Fryxell 7m 24.2 6.10 673 5

Lake Fryxell 8m 39.6 9.18 1,010 6

138

Lake Fryxell 9m 61.5 12.0 1,200 11

Lake Fryxell 10m 63.1 13.1 1,310 14

9 Lake Fryxell 11m 73.7 14.5 1,420 20

Lake Fryxell 12m 69.7 13.4 1,300 18

Lake Fryxell 15m 92.7 17.1 1,680 27

Lake Fryxell 18m 99.1 19.6 1,770 33

139

Li Rb Sr Ba McMurdo Dry Valleys Lakes µg/l µg/l µg/l µg/l

E. Lake Bonney 6m 29.3 2.80 624 9

E. Lake Bonney 22m 4,900 150 32,500 150

E. Lake Bonney 35m 8,290 294 45,100 140

W. Lake Bonney 5m 22.4 1.50 399 4

W. Lake Bonney 17m 2,810 96.1 36,500 240

0 W. Lake Bonney 37m 4,400 156 50,400 200

Lake Vanda 10m 114 2.50 308 8

Lake Vanda 62m 7,740 37.0 18,600 230

Lake Vanda 70m 22,000 94.3 69,300 660

Don Juan Pond 265,000 102 1,040,000 1,400

Seawater (Pilson, 1998) 178 120 8,130 14.1

140

Appendix E: Map of the Great Basin

141

142

Appendix F: Map of the Saskatchewan Lakes

143

144 Appendix G: Major Ion Concentrations of the Lakes in This Study

145

Saskatchewan F Cl Br SO4 Na K Mg Ca alkalinity Si T pH Lakes mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg C/L mg/L °C

Ceylon L. DL 6,980 100 196,000 23,500 2,600 2,320 33 7.88 850 nd nd

Deadmoose L. 11 8,160 8 14,800 2,240 4,20 525 34 8.35 730 nd nd

Freefight L. DL 4,190 43 92,600 8,750 2,100 2,140 7.9 7.91 6,000 nd nd

Little Manitou L. DL 21,300 93 51,100 4,460 910 2,080 85 8.04 850 nd nd

Waldsea L. 1.0 8,160 DL 30,200 2,160 610 1,320 120 8.54 1,049 nd nd

Great Basin Lakes F Cl Br SO4 Na K Mg Ca alkalinity Si T pH mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg C/L mg/L °C

Mono L. 55 16,300 15 8,970 33,700 1,500 6.24 3.7 9.87 32,000 nd nd

Pyramid L. 1.5 1,840 1.7 276 459 94 31.9 1.3 9.22 1,300 nd nd

Walker L. 5m 22 3,580 3.1 3,430 6,350 260 40.3 2.2 9.52 4,400 nd nd

Abert L. 3.8 29,800 73 1,370 51,500 1,400 DL DL 9.88 37,000 nd nd

Goose L. 0.92 602 DL 307 467 96 1.31 2.5 9.70 2,900 nd nd

146

McMurdo Dry F Cl Br SO4 Na K Mg Ca alkalinity Si T pH Valleys Lakes mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg C/L mg/L °C

Lake Hoare 4m 0.29 33.5 DL 13.5 20.3 3.6 3.64 13 9.05 10.4 1.2 0.00

Lake Hoare 5m 0.26 33.7 DL 17.9 24.0 3.8 4.55 13 8.92 10.3 1.1 0.32

Lake Hoare 6m 0.63 73.6 DL 38.6 54.1 8.9 10.1 27 8.85 22.6 2.2 0.57

Lake Hoare 8m 1.1 116 DL 62.3 90.6 15.0 16.8 40 8.70 36.6 3.4 0.40

Lake Hoare 10m 1.3 147 DL 77.2 112 18.7 21.1 49 8.42 47.6 5.1 0.47

Lake Hoare 12m 1.3 152 DL 79.4 120 20.0 22.3 51 8.14 51.4 5.3 0.55

Lake Hoare 14m 1.5 176 DL 90.3 137 22.8 25.0 58 7.85 60.1 5.8 0.45

Lake Hoare 16m 1.6 186 DL 95.1 145 24.0 26.8 64 7.64 67.2 6.3 0.26

Lake Hoare 18m 1.7 195 DL 99.3 154 25.4 28.3 67 7.54 73.5 6.7 0.24

Lake Hoare 20m 1.7 200 DL 102 156 25.9 28.9 68 7.46 76.1 6.6 0.24

Lake Hoare 22m 1.7 200 DL 101 157 25.7 28.6 68 7.37 77.5 6.7 0.24

Lake Hoare 25m 1.8 207 DL 105 160 26.6 29.6 70 7.34 79.4 6.4 0.20

Lake Hoare 30m 1.7 206 DL 104 165 27.4 30.4 72 7.28 86.5 7.0 0.17

147

McMurdo Dry F Cl Br SO4 Na K Mg Ca alkalinity Si T pH Valleys Lakes mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg C/L mg/L °C

Lake Fryxell 6m 1.1 416 DL 62.1 296 29.4 34.8 59 7.95 96 4.4 -0.03

Lake Fryxell 7m 1.2 429 DL 63.4 311 30.6 37.5 61 7.95 130 4.7 0.44

Lake Fryxell 8m 1.7 728 DL 102 532 47.8 61.5 89 7.86 164 6.8 1.33

Lake Fryxell 9m 2.0 1,090 DL 124 850 68.6 97.9 94 7.54 259 9.4 1.55

1 48 Lake Fryxell 10m 2.4 1,330 DL 144 980 77.8 110 93 7.50 296 10.6 1.69

Lake Fryxell 11m 3.8 1,710 DL 156 1,300 101 146 100 7.48 403 12.8 1.84

Lake Fryxell 12m 3.2 1,850 DL 152 1,330 104 158 100 7.53 407 12.1 1.91

Lake Fryxell 15m 5.3 2,890 DL 163 2,160 158 248 130 7.50 591 17.3 2.06

Lake Fryxell 18m 5.5 3,680 DL 105 2,680 186 319 130 7.46 694 16.8 2.14

148

F Cl Br SO4 Na K Mg Ca alkalinity Si T McMurdo Dry Valleys Lakes pH mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg/l mg C/L mg/L °C

E. Lake Bonney 6m DL 679 DL 158 354 13.1 51.3 72 8.71 19.4 nd 1.57

E. Lake Bonney 22m DL 111,000 DL 3,010 27,700 1,480 21,400 1,400 6.05 230 nd 4.69

E. Lake Bonney 35m DL 160,000 DL 3,170 48,500 2,780 31,500 1,600 6.76 54.1 nd 0.96

W. Lake Bonney 5m DL 548 DL 124 286 9.27 39 59 8.75 16.2 nd 0.85

W. Lake Bonney 17m DL 51,300 DL 3,750 23,300 701 5,360 1,800 6.07 513 nd 0.14

1 49 W. Lake Bonney 37m DL 86,800 DL 4,600 39,700 1,290 9,180 2,300 5.74 771 nd -3.91

Lake Vanda 10m DL 251 DL 16.9 52.7 12.2 19.5 81 8.35 10.3 nd 4.20

Lake Vanda 62m DL 20,400 DL 194 1,710 152 2,000 6,900 6.72 29.7 nd 15.03

Lake Vanda 70m DL 58,900 DL 525 5,000 451 6,470 21,000 5.91 68.8 nd 21.09

Don Juan Pond 12 283,000 DL 142 12,200 150 2,000 140,000 6.15 nd nd nd

Seawater (Pilson, 1998) 1.3 19,830 69.0 2,779 11,048 409 1,316 422.1 8.14** 129*

149

Appendix H: Saturation Indices for the Minerals and Lakes Discussed in Text

150

·10H

∙2H

3 3

4 4

°C CaSO CaCO BaSO CaCO SrSO CaSO MgCO Na SrCO BaCO

Barite

Calcite

Gypsum

Celestite

Witherite

Aragonite

Mirabilite

Anhydrite

Magnesite Strontianite

Log Ksp -4.362 -8.336 -9.97 -8.406 -6.630 -4.581 -7.834 -1.214 -9.271 -8.652

1 20 -1.30 -1.03 0.09 -0.83 -0.30 -1.08 1.22 -0.73 -1.24 -4.97 Ceylon Lake 0 -1.31 -1.43 0.59 -1.22 -0.17 -1.11 0.96 0.25 -1.29 -4.98

Deadmoose 20 -1.44 -0.08 0.28 0.12 -1.32 -1.20 1.25 -2.74 -1.17 -3.69 Lake 0 -1.33 -0.37 0.73 -0.17 -1.23 -1.10 0.97 -1.69 -1.28 -3.76

Freefight 20 -2.01 -0.67 0.02 -0.48 -2.35 -1.77 2.02 -1.54 -2.24 -3.98 Lake 0 -1.95 -1.03 0.50 -0.82 -2.24 -1.74 1.74 -0.52 -2.32 -4.01 Little 20 -0.93 -0.09 0.62 0.10 -0.62 -0.70 1.57 -2.12 -0.99 -3.88 Manitou Lake 0 -0.85 -0.41 1.07 -0.21 -0.52 -0.64 1.29 -1.09 -1.09 -3.93 Waldsea 20 -0.84 0.63 0.55 0.82 -0.71 -0.6 1.85 -2.71 -0.46 -3.31 Lake 0 -0.75 0.33 1.01 0.54 -0.62 -0.52 1.59 -1.66 -0.54 -3.36

151

·10H

∙2H

3 3

4 4

°C CaSO CaCO BaSO CaCO SrSO CaSO MgCO Na SrCO BaCO

Barite

Calcite

Gypsum

Celestite

Witherite

Aragonite

Mirabilite

Anhydrite

Magnesite Strontianite

52 Mono 20 -6.49 0.96 -3.03 1.15 ND -6.24 1.20 -10.11 ND -0.92

Lake 0 -6.25 0.94 -2.40 1.14 ND -6.03 1.23 -8.84 ND -0.66

Pyramid 20 -3.73 0.11 -0.01 0.30 -2.71 -3.48 1.45 -5.14 -0.09 -1.50 Lake 0 -3.54 -0.09 0.49 0.12 -2.59 -3.31 1.23 -4.08 -0.14 -1.50

Walker 20 -5.99 0.66 -1.73 0.85 -3.51 -5.74 1.92 -11.11 1.92 -0.42 Lake 5m 0 -5.71 0.57 -1.13 0.78 -3.27 -5.48 1.84 -9.99 2.01 -0.29

Abert 20 ND ND -4.03 ND ND ND ND -10.73 ND -0.93 Lake 0 ND ND -3.41 ND ND ND ND -9.48 ND -0.67

Goose 20 -3.79 0.71 -0.90 0.90 -2.94 -3.54 0.43 -5.16 0.35 -1.73 Lake 0 -3.53 0.62 -0.32 0.83 -2.71 -3.30 0.35 -4.04 0.44 -1.61

152

·10H

∙2H

3 3

4 4

CaSO CaCO BaSO CaCO SrSO CaSO MgCO Na SrCO BaCO

Barite

Calcite

Gypsum

Celestite

Witherite

Aragonite

Mirabilite

Anhydrite

Magnesite Strontianite

Lake Hoare 4m -3.16 -0.89 -1.40 -0.68 -3.18 -2.93 -1.58 -7.63 -1.92 -4.57 Lake Hoare 5m -2.85 -0.75 -1.10 -0.54 -2.83 -2.62 -1.36 -7.07 -1.74 -4.44

Lake Hoare 6m -2.44 -0.44 -0.67 -0.24 -2.41 -2.20 -1.00 -6.31 -1.43 -4.12 Lake Hoare 8m -2.22 -0.39 -0.44 -0.18 -2.17 -1.99 -0.90 -5.88 -1.35 -4.05 Lake Hoare 10m -2.09 -0.61 -0.31 -0.41 -2.03 -1.85 -1.11 -5.61 -1.57 -4.28 Lake Hoare 12m -2.03 -0.68 -0.26 -0.47 -1.98 -1.80 -1.17 -5.51 -1.65 -4.36 Lake Hoare 14m -1.99 -0.92 -0.22 -0.72 -1.95 -1.75 -1.42 -5.42 -1.90 -4.60 Lake Hoare 16m -1.90 -1.04 -0.13 -0.83 -1.87 -1.67 -1.54 -5.27 -2.02 -4.71 Lake Hoare 18m -1.88 -1.18 -0.12 -0.98 -1.86 -1.65 -1.68 -5.25 -2.17 -4.87 Lake Hoare 20m -1.87 -1.22 -0.12 -1.01 -1.86 -1.64 -1.71 -5.22 -2.21 -4.91 Lake Hoare 22m -1.88 -1.26 -0.11 -1.06 -1.85 -1.64 -1.76 -5.21 -2.25 -4.95 Lake Hoare 25m -1.86 -1.32 -0.11 -1.12 -1.85 -1.63 -1.83 -5.20 -2.32 -5.02 Lake Hoare 30m -1.85 -1.40 -0.09 -1.19 -1.84 -1.62 -1.90 -5.17 -2.39 -5.08

153

·10H

∙2H

3 3

4 4

CaSO CaCO BaSO CaCO SrSO CaSO MgCO Na SrCO BaCO

Barite

Calcite

Gypsum

Celestite

Witherite

Aragonite

Mirabilite

Anhydrite

Magnesite Strontianite

Lake Fryxell 6m -2.66 -0.55 -1.19 -0.34 -2.70 -2.43 -0.92 -5.39 -1.59 -4.53 Lake Fryxell 7m -2.65 -0.40 -1.21 -0.19 -2.70 -2.42 -0.75 -5.38 -1.46 -4.40 Lake Fryxell 8m -2.42 -0.27 -1.08 -0.07 -2.45 -2.18 -0.57 -4.86 -1.34 -4.38 Lake Fryxell 9m -2.42 -0.44 -0.85 -0.24 -2.40 -2.18 -0.57 -4.46 -1.46 -4.32 Lake Fryxell 10m -2.40 -0.46 -0.72 -0.25 -2.34 -2.16 -0.52 -4.31 -1.43 -4.22 Lake Fryxell 11m -2.41 -0.35 -0.62 -0.14 -2.35 -2.16 -0.33 -4.10 -1.33 -4.00 Lake Fryxell 12m -2.43 -0.29 -0.69 -0.09 -2.41 -2.19 -0.24 -4.11 -1.32 -4.00 Lake Fryxell 15m -2.46 -0.15 -0.64 0.06 -2.43 -2.22 0.01 -3.79 -1.16 -3.78 Lake Fryxell 18m -2.71 -0.14 -0.83 0.07 -2.67 -2.47 0.11 -3.87 -1.14 -3.70

154

·10H

∙2H

3 3

4 4

CaSO CaCO BaSO CaCO SrSO CaSO MgCO Na SrCO BaCO

Barite

Calcite

Gypsum

Celestite

Witherite

Aragonite

Mirabilite

Anhydrite

Magnesite Strontianite

Lake Vanda 10m -2.57 -0.82 -1.12 -0.62 -3.05 -2.31 -1.58 -7.15 -2.38 -4.83

Lake Vanda 60m -1.42 -0.62 -0.25 -0.42 -1.69 -1.15 -1.30 -5.37 -2.07 -4.87 Lake Vanda 70m -0.60 -0.75 -0.09 -0.56 -1.17 -0.40 -1.34 -4.43 -2.54 -5.56 East Lake Bonney 6m -1.82 -0.44 -0.25 -0.23 -1.96 -1.58 -0.70 -4.51 -1.62 -4.32 East Lake Bonney 22m -0.58 -0.89 0.59 -0.68 -0.42 -0.47 0.37 -1.42 -1.76 -5.16 East Lake Bonney 35m -0.26 -0.25 0.49 -0.04 -0.06 -0.28 1.25 -1.69 -1.14 -4.94 West Lake Bonney 5m -1.94 -0.47 -0.56 -0.27 -2.19 -1.70 -0.77 -4.72 -1.75 -4.54 West Lake Bonney -0.53 -1.01 0.99 -0.80 -0.30 -0.33 -0.59 -1.08 -1.79 -4.93 15m West Lake Bonney -0.52 -1.07 0.63 -0.86 -0.36 -0.36 -0.49 -0.81 -1.93 -5.36 35m Don Juan Pond -0.37 ND -1.69 ND -1.25 -1.26 ND -9.61 ND ND

155

References

Aiuppa, A., Allard, P., D’Alessandro, W., Michel, A., Parello, F., Treuil, M., and Valenza, M., 2000. Mobility and fluxes of major, minor and trace metals during basalt weathering and groundwater transport at Mt. Etna volcano (Sicily). Geochim. Cosmochim. Acta 64, 1827- 1841.

Amphlett, C., 1964. Inorganic Ion Exchangers. Elsevier.

Anderson, M., Bertsch, P., and Miller, W., 1989. Exchange and apparent fixation of lithium in selected soils and clay minerals. Soil Science 148, 46-52.

Anghel, I., Turin, H., and Reimus, P., 2002. Lithium sorption to Yucca Mountain tuffs. Applied Geochemistry 17, 819-924.

Angino, E., Armitage, K., and Tash, J., 1962. Chemical stratification in Lake Fryxell, Victoria Land, Antarctica. Science 138, 34-36.

Angino, E., Armitage, K., and Tash, J., 1965. A chemical and limnological study of Lake Vanda, Victoria Land, Antarctica. University of Kansas Science Bulletin 45, 1097-1118.

Bein, A. and Arad, A., 1992. Formation of saline groundwaters in the Baltic region during glacial periods. Journal of Hydrology 140, 75-87.

Benson, L., 1994. Carbonate deposition, Pyramid Lake subbasin, Nevada. 1. Sequence of formation and elevational distribution of carbonate deposits (tufas). Palaeogeography Palaelclimatology Palaeoecology 109, 55-87.

Benson, L., Braddock, L., Smoot, J., Mensing, S., Lund, S., Stine, S., and Sarna-Wojciki, A., 2003. Influence of the Pacific Decadal Oscillation on the climate of the Sierra Nevada, California and Nevada. Quaternary Research 59, 151-159.

Benson, L., Kashgarian, M., Rye, R., Lund, S., Paillet, F., Smoot, J., Kester, C., Mensing, S., Meko, D., and Lindstrom, S., 2002a. Holocene multidecadal and multicentennial droughts affecting Northern California and Nevada. Quaternary Science Reviews 21, 659-682.

Benson, L. and Leach, D., 1979. Uranium transport in the Walker River basin, California and Nevada. Journal of Geochemical Exploration 11, 227-248.

156

Benson, L., Meyers, P., and Spencer, R., 1991. Change in the size of Walker Lake during the past 5000 years. Palaeogeography Palaelclimatology Palaeoecology 81, 189-214.

Benson, L. and Paillet, F., 2002b. HIBAL: a hydrologic-isotopic-balance model for application to paleolake systems. Quaternary Science Reviews 21, 1521-1539.

Benson, L. and Spencer, R., 1983. A hydrological reconnaissance study of the Walker River Basin, California and Nevada.

Benson, L. and White, J., 1994. Stable Isotopes of Oxygen and Hydrogen in the Truckee River- Pyramid Lake Surface-Water System. 3. Source of Water Vapor Overlying Pyramid Lake. Limnology and Oceanography 39, 1945-1958.

Bernstein, R. and Byrne, R., 2004. Acantharians and marine barite. Marine Chamistry 86, 45-50.

Bernstein, R., Byrne, R., PR, B., and Greco, A., 1992. Morphologies and transformations of celestite in seawater: the role of acantharians in strontium and barium geochemistry. Geochim. Cosmochim. Acta 56, 3272-3279.

Bertler, N. A. N., Mayewski, P., Barrett, P., Sneed, S. B., Handley, M. J., and Kreutz, K., 2004b. Monsoonal circulation of the McMurdo Dry Valleys--signal from snow chemistry. Annals of Glaciology 39, 139-145.

Betcher, R., Grove, G., and Pupp, C., 1995. Groundwater in Manitoba: Hydrogeology, quality concerns, management. National Hydrology Institute.

Bischoff, J., Fitzpatrick, J., and Rosenbauer, R., 1993a. The solubility and stabilization of ikaite (CaCO3·6H2O) from 0° to 25°C: Environmental and paleoclimatic implications for thinolite tufa. The Journal of Geology 101, 21-33.

Bischoff, J., Stine, S., Rosenbauer, R., Fitzpatrick, J., and Stafford, T. J., 1993b. Ikaite precipitation by mixing of shoreline springs and lake water, Mono Lake, California, USA. Geochim. Cosmochim. Acta 57, 3855-3865.

Black, R., Jackson, M., and Berg, T., 1965. Saline discharge from Taylor Glacier, Victoria Land, Antarctica. Journal of Geology 74, 175-181.

Bockheim, J., 1997. Properties and classification of cold desert soils from Antarctica. Soil Science Society of America Journal 61, 224-231.

Bockheim, J., 2002. Landform and soil development in the McMurdo Dry Valleys: a regional synthesis. Arctic, Antarctic, and Alpine Research 34, 308-317.

Bottomley, D., Chan, L.-H., Katz, A., Starinsky, A., and Clark, I., 2003. Lithium isotope geochemistry and origin of Canadian shield brines. Ground Water 41, 847-856.

157

Bottomley, D., Katz, A., Chan, L.-H., Starinsky, A., Douglas, M., Clark, I., and Raven, K., 1999. The origin and evolution of Canadian Shield brines: evaporation or freezing of seawater? New lithium isotope and geochemical evidence from the Slave craton. Chemical Geology 155, 295-320.

Broecker, W., 1971. A kinetic model for the chemical composition of seawater. Quaternary Research 1, 188-207.

Broecker, W. and Peng, T.-H., 1982. Tracers in the Sea. Lamont-Doherty Geological Observatory, Palisades, N.Y.

Buetel, M., Horne, A., Roth, J. C., and Barratt, N. J., 2001. Limnological effects of anthropogenic dessication of a large, saline lake, Walker Lake, Nevada. Hydrobiologia 466, 91-105.

Campbell, I. and Claridge, G., 1982. The influence of moisture on the development of soils of the cold deserts of Antarctica. In: Yaalon, D. (Ed.), Aridic soils. Elsevier, Amsterdam.

Campbell, I. and Claridge, G., 1987. Antarctica: Soils, Weathering Processes and Environment. Elsevier, Amsterdam.

Campbell, S. and Wilson, A. T., 1972. Bathymetry and bottom sediments of Lake Vanda, Antarctica. Antarctic Journal of the United States July-August, 97-99.

Carlson, C., Phillips, F., Elmore, D., and Bentley, H., 1990. Chlorine 36 tracing of salinity sources in the Dry Valleys of Victoria Land, Antarctica. Geochim. Cosmochim. Acta 54, 311-348.

Chan, L.-H., Drummond, D., Edmond, J. M., and Grant, B., 1977. On the barium data from the Atlantic GEOSECS Expedition. Deep-Sea Research 24, 613-649.

Chan, L.-H. and Edmond, J. M., 1988. Variation of lithium isotope composition in the marine environment: A preliminary report. Geochim. Cosmochim. Acta 52, 1711-1717.

Charlton, S. and Parkhurst, D., 2002. PHREEQCI; a graphical user interface to the geochemical model PHREEQC. In: Survey, F. S.-U. S. G. (Ed.). U. S. Geological Survey, Denver, CO.

Chinn, T., 1981. Hydrology and Climate in the Ross Sea Area. Journal of the Royal Society of New Zealand 11, 373-386.

Chinn, T., 1990. The dry valleys. In: Hatherton, T. (Ed.), Antarctica; the Ross Sea region. Department of Scientific and Industrial Research (DSIR), Wellington.

Chinn, T., 1993. Physical hydrology of the Dry Valley lakes. Antarctic Research Series 59, 1-51.

Claridge, G. and Campbell, I., 1977. The salts in Antarctic soils, their distribution and relationship to soil processes. Soil Science 123, 377-384.

158

Clow, G., McKay, C., Simmons, G. J., and Wharton, R., 1988. Climatological observations and predicted sublimation rates at Lake Hoare, Antarctica. Journal of Climate 1, 715-728.

Conovitz, P., MacDonald, L., and McKnight, D. M., 2006. Spatial and temporal active layer dynamics along three glacial meltwater streams in the McMurdo Dry valleys, Antarctica. Arctic, Antarctic, and Alpine Research 38, 42-53.

Conovitz, P., McKnight, D., MacDonald, L., Fountain, A., and House, H., 1998. Hydrologic processes influencing streamflow variation in Fryxell Basin, Antarctica. In: Priscu, J. (Ed.), Ecosystem Dynamics in a Polar Desert; the McMurdo Dry Valleys, Antarctica. American Geophysical Union, Washington, DC.

Das, A. and Krishnaswamin, S., 2006. Barium in Deccan basalt rivers: Its abundance, relative mobility and flux. Aq. Geochem. 12, 221-238. de Lange, G., Boelrijk, N., Catalano, G., Cordelli, C., Klinkhammer, G. P., Middelburg, J., Muller, D., Ullman, W., van Gaans, P., and Woittiez, J., 1990b. Sulphate-related equilibria in the hypersaline brines of the Tyro and Bannock Basins, eastern Mediterranean Marine Chemistry 31, 89-112.

DeHairs, F., Chesslet, R., and Jedwab, J., 1980. Discrete suspended particles of barite and the barium cycle in the open ocean. Earth and Planetary Science Letters 49, 528-550.

DeHairs, F., Lambert, C. E., Chesslet, R., and Risler, N., 1987. The biological production of marine suspended barite and the barium cycle in the Western Mediterrean Sea. Biogeochemistry 4, 119-139.

Dieckmann, G., Nahrke, G., Papadimitriou, S., Gottlicher, J., R, S., Kennedy, H., Wolf-Gladrow, D., and Thomas, D., 2008. Calcium carbonate as ikaite crystals in Antarctic sea ice. Geophys. Res. Lett. 35, L08501.

Domagalski, J., Orem, W., and Eugster, H., 1989. Organic geochemistry and brine composition in Great Salt, Mono, and Walker Lakes. Geochim. Cosmochim. Acta 53, 2857-2872.

Doran, P., Clow, G., Fountain, A., Fritsen, C., Lyons, W., McKay, C., McKnight, D., Moorhead, D., Parsons, A., Priscu, J., Virginia, R., Wall, D., and Walsh, J., 2002. Antarctic climate cooling and terrestrial ecosystem response. Nature 415, 517-520.

Doran, P., McKay, C., Fountain, A., Nylen, T. H., McKnight, D. M., Jaros, C., and Barrett, J. E., 2008. Hydrologic response to extreme warm and cold summers in the McMurdo Dry Valleys, East Antarctica. Ant. Sci. 20, 499-509.

Doran, P., Wharton, R., and Lyons, W. B., 1994. Paleolimnology of the McMurdo Dry Valleys, Antarctica. Journal of Paleolimnology 10, 85-114.

Drever, J. I., 1988. The Geochemistry of Natural Waters. Prentice Hall, Englewood Cliffs, NJ. 159

Edmond, J. M., Measures, C., McDuff, R., Chan, L.-H., Collier, R. W., Grant, B., Gordon, L., and Corliss, J., 1979. Ridge crest hydrothermal activity and the balances of major and minor elements in the ocean: The Galapogos data. Earth and Planetary Science Letters 46, 1-18.

Edmond, J. M., Palmer, M. R., Measures, C., Brown, E., and Huh, Y., 1996. Fluvial geochemistry of the eastern slope of the northeastern Ander and its foredeep in the drainage of the Orinoco in Columbia and Venezuela. Geochim. Cosmochim. Acta 60, 2949-2975.

Edmond, J. M., Palmer, M. R., Measures, C., Grant, B., and Stallard, R., 1995. The fluvial geochemistry and dunedation rate of the Guyana shield in Venezuela, Columbia, and Brazil. Geochim. Cosmochim. Acta 59, 3301-3325.

Elderfield, H. and Gieskes, J., 1982. Sr isotopes in interstitial waters of marine sediments from Deep Sea Drilling Project cores. Nature 300, 493-496.

Elderfield, H. and Schultz, A., 2006. Mid-ocean ridge hydrothermal fluxes and the chemical composition of the ocean. Annual Reviews of Earth and Planetary Sciences 24, 191-224.

Elston, D. and Bressler, S., 1981. Magnetic stratigraphy of DVDP drill cores and Late Cenozoic history of Taylor Valley, Transantarctic Mountains, Antarctica, Dry Valley Drill Project. American Geophysical Union.

Eugster, H. and Hardie, L., 1978. Saline Lakes. In: Lerman, A. (Ed.), Lakes: Chemistry, Geology, Physics. Springer-Verlag, New York, NY.

Eugster, H. and Jones, B. F., 1979. Behavior of major solutes during closed-basin brine evolution. American Journal of Science 279, 609-631.

Falkner, K. K., Klinkhammer, G. P., Bowers, T. S., Todd, J. E., Lewis, B. L., Landing, W. M., and Edmond, J. M., 1993. The behavior of barium in anoxic marine waters. Geochim. Cosmochim. Acta 57, 537-554.

Faure, G. and Mensing, T., 2005. Isotopes: Principles and Applications. John Wiley and Sons, Inc., Hoboken, N.J.

Fleming, T., Foland, K., and Elliot, D., 1995. Isotopic and chemical constraints on the crustal evolution and source signature of Ferrar magmas, north Victoria Land, Antarctica. Contributions to Mineralolgy and Petrology 121, 217-236.

Fortner, S., Lyons, W. B., Munk, L., and Cozzetto, K., in review. Landscape and flow controls of Antarctic dry valley glacier melt stream geochemistry Water Resources Research.

Fortner, S., Tranter, M., Fountain, A., Lyons, W. B., and Welch, K., 2005. The geochemistry of supraglacial streams of Canada Glacier, Taylor Valley (Antarctica), and their evolution into proglacial waters. Aq. Geochem. 11, 391-412.

160

Fountain, A., Lyons, W., Burkins, M., Dana, G., Doran, P., Lewis, K., McKnight, D., Moorhead, D., Parsons, A., Priscu, J., Wall, D., Warton, R., and Virginia, R., 1999b. Physical controls on the Taylor Valley ecosystem, Antarctica. Bioscience 49, 961-971.

Frape, S., Fritz, P., and McNutt, R., 1984. The role of water-rock interaction in the Canadian Shield. Geochim. Cosmochim. Acta 48, 1617-1627.

Fritz, S., 1992. Measuring the ratio of aqueous diffusion coefficients between 6Li+Cl- and 7Li+Cl- by osmometry. Geochim. Cosmochim. Acta 56, 3781-3789.

Froelich, P., 1980. Analysis of organic carbon in marine sediments. Limnology and Oceanography 25, 564-572.

Gac, J., Droubi, A., Fritz, B., and Tardy, Y., 1977. Geochemical behavior of silica and magnesium during evaporation of waters in Lake Chad. Chemical Geology 19, 215-228.

Gaillardet, J., Viers, J., and Dupré, B., 2003. Trace Elements in River Waters. In: Drever, J. I. (Ed.), Treatise on Geochemistry. Elsevier.

Galat, D. and Jacobson, R., 1985. Recurrent aragonite precipitation in saline-alkaline Pyramid Lake, Nevada. Archives of Hydrobiology Supplement 105, 137-159.

Galat, D., Lider, E., Vigg, S., and Robertson, S., 1981. Limnoloy of a large, deep, North American terminal lake, Pyramid Lake, Nevada, U.S.A. Hydrobiologia 81-2, 281-317.

Garrels, R. and Mackenzie, F., 1967. Origin of the chemical composition of some springs and lakes, Equilibrium Concepts in Natural Water Systems. American Chemical Society.

Gislason, S. R., Arnorsson, S., and Armannsson, H., 1996. Chemical weathering of basalt as deduced from the composition of precipitation, rivers and rocks in SW Iceland. American Journal of Science 296, 837-907.

Gislason, S. R. and Eugster, H. P., 1987. Meteoric water-basalt interactions. II: A field study in NE Iceland Geochim. Cosmochim. Acta 51.

Gooseff, M., McKnight, D., Lyons, W., and Blum, A., 2002. Weathering reactions and hyporheic exchange controls on stream water chemistry in a glacial meltwater stream in the McMurdo Dry Valleys. Water Resources Research 38, 1279-1294.

Gooseff, M., McKnight, D. M., and Runkel, R., 2004. Reach-scale cation exchange controls on major ion chemistry of an Antarctic glacial meltwater stream. Aq. Geochem. 10, 221-238.

Green, W., Angle, M., and Chave, K., 1988. The geochemistry of Antarctic streams and their role in the evolution of four lakes of the McMurdo dry valleys. Geochim. Cosmochim. Acta 52, 1265-1274.

161

Green, W., Stage, B., Preston, A., Wagers, S., Shacat, J., and Newell, S., 2005. Geochemical processes in the Onyx River, Wright Valley, Antarctica: Major ions, nutrients, trace metals. Geochim. Cosmochim. Acta 69, 839-850.

Green, W. J. and Canfield, D., 1984. Geochemistry of the Onyx River (Wright Valley, Antarctica) and its role in the chemical evolution of Lake Vanda. Geochim. Cosmochim. Acta 48, 2457- 2467.

Green, W. J., Gardner, T. J., Ferdelman, T. G., Angle, M., Varner, L., and Nixson, P., 1989. Geochemical processes in the Lake Fryxell basin (Victoria Land Antarctica). Hydrobiologia 172, 129-148.

Green, W. J. and Lyons, W. B., 2009. The saline lakes of the McMurdo Dry Valleys, Antarctica. Aq. Geochem. 15, 321-348.

Hall, B. L. and Denton, G., 2005. Surficial geology and geomorphology of eastern and central Wright Valley, Antarctica. Geomorphology 64, 25-65.

Hardie, L. and Eugster, H., 1970. The evolution of closed-basin brines, Fiftieth Anniversary Symposia, Mineralogy and Geochemistry of Non-Marine Evaporites.

Harris, H. J. H. and Cartwright, K., 1981. Hydrology of the Don Juan Pond basin, Wright Valley, Antarctica. Antarctic Research Series 33, 161-184.

Haskell, T., Kennett, J., Prebble, W., Smith, G., and Willis, I., 1965. The geology of the middle and lower Taylor Valley of South Victoria Land, Antarctica. Transactions of the Royal Society of New Zealand--Geology 2, 169-186.

Henderson, G., Hall, B. L., Smith, A., and Robinson, L. F., 2006. Control on (234U/238U) in lake water: A study in the Dry Valleys of Antarctica. Chemical Geology 226, 298-308.

Hendy, C. H., Healy, T. R., Rayner, E. M., Shaw, J., and Wilson, A. T., 1979. Late pleistocene glacial chronology of the Taylor Valley, Antarctica, and the global climate. Quaternary Research 11, 172-184.

Hinkley, T. K. and Matsumoto, A., 2001. Atmospheric regime of dust and salt through 75,000 years of Taylor Dome ice core: Refinement by measurement of major, minor, and trace metal suites. Journal of Geophysical Research 106, 18487-18493.

Hosler, W., 1979. Trace elements and isotopes in evaporites. In: Burns, R. (Ed.),Marine Minerals Reviews in Mineralology. Mineralogical Society of America, Washington DC.

Howard-Williams, C., Hawes, I., and Schwarz, A., 1997. Sources and sinks of nutrients in a polar desert stream, the Onyx River, Antarctica. In: Lyons, W. B., Howard-Williams, C., and Hawes, I. Eds.), Ecosystem Processes in Antarctic Ice-Free Landscapes. Balkema.

162

Huh, Y., Chan, L.-H., and Edmond, J., 2001. Lithium isotopes as a probe of weathering processes; Orinoco River. Earth and Planetary Science Letters 194, 189-199.

Huh, Y., Chan, L.-H., Zhang, L., and Edmond, J., 1998. Lithium and its isotopes in major world rivers: implications for weathering and the oceanic budget. Geochim. Cosmochim. Acta 62, 2039-2051.

Ishman, S. and Rieck, H., 1992. A late Neogene Antarctic glacioeustatic record, Victoria Land basin margin, Antarctica, The Antarctic Paleoenvironment: A Perspective on Global Change. American Geophysical Union.

James, R. H. and Palmer, M. R., 2000. Marine geochemical cycles of the alkali elements and boron: The role of sediments. Geochim. Cosmochim. Acta 64, 3111-3122.

Jehl, J., 1983. Tufa formation at Mono Lake, California. California Geology 36, 3.

Jellison, R. and Melack, J., 1993. Meromixis in Hypersaline Mono Lake, California. 1. Stratification and Vertical Mixing During the Onset, Persistence, and Breakdown of Meromixis Limnology and Oceanography 38, 1008-1019.

Jones, B. F., 1966. Geochemical evolution of closed-basin water in the western Great BasinProceedings of the Second Symposium on Salt.

Jones, B. F. and Deocampo, D. M., 2004. Geochemistry of Saline Lakes. In: Drever, J. I. (Ed.), Surface, Groundwater, Weathering, and Soils. Elsevier-Pergamon, Oxford, UK and San Diego, USA.

Jones, L. and Faure, G., 1967. Origin of the salts in Lake Vanda, Wright Valley, southern Victoria Land, Antarctica. Earth and Planetary Science Letters 3, 101-106.

Jones, L. and Faure, G., 1978. A study of strontium isotopes in lakes and surficial deposits of the ice-free valleys, southern Victoria Land, Antarctica. Chemical Geology 22, 107-120.

Keller, K., Blum, J., and Kling, G., 2007. Geochemistry of soils and streams on surfaces of varying ages in arctic Alaska. Arctic, Antarctic, and Alpine Research 39, 84-98.

Keys, J. and Williams, K., 1981. Origin of crystalline, cold desert salts in the McMurdo region, Antarctica. Geochim. Cosmochim. Acta 45, 2299-2309.

Krumgalz, B., Starinsky, A., and Pitzer, K., 1999. Ion-interaction approach: Pressure effect on the solubility of some minerals in submarine brines and seawater. Journal of Solution Chemistry 28, 667-692.

Lancaster, N., 2002. Flux of eolian sediment in the McMurdo Dry Valleys, Antarctica: A preliminary assessment. Arctic, Antarctic, and Alpine Research 34, 318-323.

163

Last, W. M., 1989. Continental brines and evaporites of the northern Great Plains of Canada. Sedimentary Geology 64, 207-221.

Last, W. M., 1992. Chemical composition of saline and subsaline lakes of the northern Great Plains, western Canada. International Journal of Salt Lake Research 1, 47-76.

Last, W. M. and Ginn, F. M., 2005. Saline systems of the Great Plains of western Canada: an overview of the limnogeology and paleolimnology. Saline Systems 1.

Last, W. M. and Schweyen, T. H., 1983. Sedimentology and geochemistry of saline lakes of the northern Great Plains. Hydrobiologia 105, 245-263.

Last, W. M. and Schweyen, T. H., 1985. Late Holocenehistory of Waldsea Lake, Saskatchewan, Canada. Quaternary Research 24, 219-234.

Last, W. M. and Slezak, L., 1986. Paleohydrology sedimentology, and geochemistry of two meromictic saline lakes in southern Saskatchewan. Geographie Physique et Quaternaire XL, 5-15.

Last, W. M. and Slezak, L., 1988. The salt lakes of western Canada: A plaeolimnological overview. Hydrobiologia 158, 301-316.

Legrand, M. and Mayewski, P., 1997. Glaciochemistry of polar ice cores: A review. Reviews of Geophysics 35, 219-243.

LeMarchand, E., Reynolds, B., Wiederhold, J., Bourdon, B., and Kretzschmar, R., 2008. Lithium isotope fractionation along a weathering sequence. Geochim. Cosmochim. Acta 72, A533.

Li, Y. and Gregory, S., 1974. Diffusion if ions in sea water and deep-sea sediments. Geochim. Cosmochim. Acta 38, 703-714.

Lopes, T. and Smith, J. L., 2007. Bathymetry of Walker Lake, West-Central NevadaScientific Investigations Report. U. S. Department of the Interior, U. S. Geological Survey.

Lyons, W., Fountain, A., Doran, P., Priscu, J., Neumann, K., and Welch, K., 2000. Importance of landscape position and legacy: the evolution of the lakes in Taylor Valley, Antarctica. Freshwater Biology 43, 355-367.

Lyons, W. and Mayewski, P., 1993. The Geochemical Evolution of Terrestrial Waters in the Antarctic: The Role of Rock-Water Interactions. Antarctic Research Series 59, 135-143.

Lyons, W., Nezat, C., Benson, L., Bullen, T., Graham, E., Kidd, J., Welch, K., and Thomas, J., 2002. Strontium isotopic signatures of the streams and lakes of Taylor Valley, southern Victoria Land, Antarctica: Chemical weathering in a polar climate. Aq. Geochem. 8, 75-95.

164

Lyons, W., Tyler, S., Wharton, R., McKnight, D., and Vaughn, B., 1998b. A Late Holocene desiccation of Lake Hoare and Lake Fryxell, McMurdo Dry Valleys, Antarctica. Ant. Sci. 10, 247-256.

Lyons, W. and Welch, K., 1997. Lithium in waters of a polar desert. Geochim. Cosmochim. Acta 61, 4309-4319.

Lyons, W., Welch, K., Fountain, A., Dana, G., Vaughn, B., and McKnight, D., 2003. Surface glaciochemistry of Taylor Valley, southern Victoria Land, Antarctica and its relationship to stream chemistry. Hydrol. Process. 17, 115-130.

Lyons, W., Welch, K., Snyder, G., Olesik, J., Graham, E., and Marion, G., 2005. Halogen geochemistry of the McMurdo dry valleys lakes, Antarctica: Clues to the origin of solutes and lake evolution. Geochim. Cosmochim. Acta 69, 305-323.

Lyons, W. B., 2009. Stream Nutrients (nitrate, nitrite, ammonium, reactive phosphorus). McMurdo Dry Valleys Long Term Ecological Research.

Lyons, W. B., Frape, S., and Welch, K., 1999b. History of the McMurdo Dry Valley lakes, Antarctica, from stable chlorine isotope data. Geology 6, 527-530.

Lyons, W. B., Welch, K., Crick, K., Toxey, J., Mastrine, J., and McKnight, D. M., 1997. Chemcial weathering rates and reactions in the Lake Fryxell Basin, Taylor Valley: Comparisono to temperate river basins. In: Lyons, W. B., Howard-Williams, C., and Hawes, I. Eds.), Ecosystem Processes in Antarctic Ice-Free Landscapes. AA Balkema, Rotterdam, Netherlands.

Lyons, W. B., Welch, K., Neumann, K., Toxey, J., McArthur, R., Williams, C., McKnight, D. M., and Moorhead, D., 1998a. Geochemical linkages among glaciers, streams, and lakes within Taylor Valley, Antarctica. In: Priscu, J. C. (Ed.), Ecosystem Dynamics in a Polar Desert: The McMurdo Dry Valleys, Antarctica. American Geophysical Union, Washington, DC.

Lyons, W. B., Welch, K., and Sharma, P., 1998c. Chlorine-36 in the waters of the McMurdo Dry Valley lakes, southern Victoria Land, Antarctica: Revisited. Geochim. Cosmochim. Acta 62, 185-191.

MacKenzie, F. and Garrels, R., 1966. Chemical mass balance between rivers and oceans. American Journal of Science 264, 507-525.

MacKenzie, F., Ristvet, B., Thorstenson, D., Lerman, A., and Leeper, R., 1981. Reverse weathering and chemical mass balance in a coastal environment. In: Martin, J.-M., Burton, J., and Eisma, D. Eds.), River Inputs to Ocean Systems. United Nations Press, New York, NY.

Marion, G., 1997. A theoretical evaluation of mineral stability in Don Juan Pond, Wright Valley, Victoria Land. Ant. Sci. 9, 92-99.

165

Marion, G. M. and Farren, R. E., 1999. Mineral solubilities in the Na-K-Mg-Ca-Cl-SO4-H2O system: A re-evaluation of the sulfate chemistry in the Spencer-Moller-Weare model. Geochim. Cosmochim. Acta 63, 1305-1318.

Marion, G. M., Farren, R. E., and Komrowski, A. J., 1999. Alternative pathways for seawater freezing. Cold Regions Science and Technology 29, 259-266.

Marriott, C., Henderson, G., Crompton, R., Staubwasser, M., and Shaw, S., 2004. Effect of mineraology, temperature, and salinity on Li/Ca and Li isotopic composition of calcium carbonate. Chemical Geology 212, 5-15.

Masuda, N., Nishimura, M., and Torii, T., 1982. Pathway and distribution of trace elements in Lake Vanda, Antarctica. Nature 298, 154-156.

Matsubaya, O., Sakai, H., Torii, T., Burton, H., and Kerry, K., 1979. Antarctic saline lakes--Stable isotope ratios, chemical compositions and evolution. Geochim. Cosmochim. Acta 43, 7-26.

Maurice, P., McKnight, D., Leff, L., Fulghum, J., and Gooseff, M., 2002. Direct observations of aluminosilicate weathering in the hyporheic zone of an Antarctic Dry Valley stream. Geochimica et Cosmochimica Acta, 66, 1335-1347.

McKnight, D., Niyogi, D., Alger, A., Bomblies, A., Conovitz, P., and Tate, C., 1999. Dry valley streams in Antarctica: ecosystems waiting for water. BioScience 49, 985–995.

McKnight, D. M., 2009. Onyx River at Vanda gauge measurements. McMurdo Dry Valleys Long Term Ecological Research.

MCM-LTER, 2007. Limnological Methods for the McMurdo Long Term Ecological Research Program.

Michalopoulos, P. and Aller, R., 1995. Rapid clay mineral formation in Amazon delta sediments: Reverse weathering and oceanic elemental cycles. Science 270, 614-617.

Michalski, G., Bockheim, J., Kendall, C., and Thiemens, M., 2005. Isotopic composition of Antarctic Dry Valley nitrate: Implications for NOy sources and cycling in Antarctica. Geophys. Res. Lett. 32, L13817.

Moore, W. S. and Falkner, K. K., 1999. Cycling of radium and barium in the Black Sea. Journal of Environmental Radioactivity 43.

Morel, F. M. M. and Herring, J. G., 1993. Principles and Applications of Aquatic Chemistry. John Wiley and Sons, Inc., New York.

Nesbitt, H., Markovics, G., and Price, R., 1980. Chemical processes affecting alkalis and alkaline earths during continantal weathering. Geochim. Cosmochim. Acta 44, 1659-1666.

166

Neumann, K., 1999. Carbon Dynamics in Lakes and Streams in Taylor Valley Antarctica, University of Alabama.

Neumann, K., Lyons, W. B., and Des Marais, D. J., 1998. Inorganic carbon-isotope distribution and budget in the Lake Hoare and Lake Fryxell basins, Taylor Valley, Antarctica. Annals of Glaciology 27, 685-690.

Nezat, C., Lyons, W., and Welch, K., 2001. Chemical weathering in streams of a polar desert (Taylor Valley, Antarctica). GSA Bulletin 113, 1401-1408.

Olsher, U., Izatt, R. M., Bradshaw, J. S., and Dalley, N. K., 1991. Coordination chemistry of lithium ion--A crystal and molecular structure review. Chemical Reviews 91, 137-164.

Omelon, C., Pollard, W., and Marion, G. M., 2001. Seasonal formation of ikaite (CaCO3·6H2O) in saline spring discharge at Expedition Fiord, Canadian High ArcticL Assessing conditional constraints for natural crystal growth. Geochim. Cosmochim. Acta 65, 1429- 1437.

Palmer, M. R. and Edmond, J. M., 1989a. The strontium isotope budget of the modern ocean. Earth and Planetary Science Letters 92, 11-26.

Parkhurst, D. and Appelo, C., 1999. PHREEQCWater-Resources Investigations Report 99-4259. U.S. Geological Survey, U.S. Department of the Interior, Denver, CO.

Phillips, K. and Van Denburgh, A., 1971. Hydrology and geochemistry of Abert, Summer, and Goose lakes, and other closed basin lakes in south-central OregonU. S. Geological Survey Professional Paper 1044-9612. U. S. Geological Survey.

Pilson, M. E. Q., 1998. An Introduction to the Chemistry of the Sea. Prentice Hall, New Jersey.

Pistiner, J. and Henderson, G., 2003. Lithium-isotope fractionation during continental weathering processes. Earth and Planetary Science Letters 215, 327-339.

Pogge von Strandmann, P., James, R. H., van Calsteren, P., Gislason, S., and Burton, K., 2008. Lithium, magnesium and uranium isotope behaviour in the estaurine environment of basaltic islands. Earth and Planetary Science Letters 274, 462-471.

Poreda, R., Hunt, A., Lyons, W. B., and Welch, K., 2004. The helium isotopic chemistry of Lake Bonney, Taylor Valley, Antarctica: Timing of Late Holocene climate change in Antarctica. Aq. Geochem. 10, 353-371.

Porter, S. and Beget, J., 1981. Provenance and depositional environments of Late Cenozoic sediments in permafrost cores from the lower Taylor Valley, Antarctica. In: McGinnis, L. (Ed.), Dry Valley Drilling Project. American Geophysical Union.

Puchelt, H., 1972. Barium. In: Wedepohl, K. M. (Ed.), Handbook of Geochemistry. Springer-Verlag. 167

Rimstidt, J., Balog, A., and Webb, J., 1998. Distribution of trace elements between carbonate minerals and aqueous solutions. Geochim. Cosmochim. Acta 62, 1851-1863.

Rosner, M., Ball, L., Peucker-Ehrenbrink, B., Blusztajn, J., Bach, W., and Erzinger, J., 2007. A simplified, accurate, and fast method for lithium isotope analysis of rocks and fluids, and 7Li values of seawater and rock reference materials. Geostandards and Geoanalytical Research 31, 77-88.

Rushdi, A. I., McManus, J., and Collier, R. W., 2000. Marine barite and celestite saturation in seawater. Marine Chemistry 69, 13-31.

Rutten, A. and De Lange, G., 2002. A novel selective extraction of barite, and its application to eastern Mediterranean sediments. Earth and Planetary Science Letters 198, 11-24.

Schenau, S., Prins, M., De Lange, G., and Monnin, C., 2001. Barium accumulation in the Arabian Sea: Controls on barite preservation in marine sediments. Geochim. Cosmochim. Acta 65, 1545-1556.

Schijf, J., 2007. Alkali elements (Na, K, Rb) and Alkaline earth elements (Mg, Ca, Sr, Ba) in the anoxic brine of Orca Basin, northern Gulf of Mexico. Chemical Geology 243, 255-274.

Schultzelam, S. and Beveridge, T., 1994. Nucleation of celestite and strontianite on a cyanobacteria s-layer. Applied and Environmental Microbiology 60, 447-453.

Seyfried, W. J., Janecky, D., and Mottl, M., 1984. Alteration of the oceanic crust: Implications for geochemical cycles of lithium and boron. Geochim. Cosmochim. Acta 48, 557-569.

Smith, G. and Friedman, I., 1993. Lithology and paleoclimatic implications of lacustrine deposits around Lake Vanda and Don Juan Pond, Antarctica. Antarctic Research Series 59, 83-94. Sonnenfeld, P., 1984. Brines and Evaporites. Academic Press, Orlando, Florida.

Spiegel, R. H. and Priscu, J. C., 1998. Physical Limnology of the McMurdo Dry Valley Lakes. In: Priscu, J. C. (Ed.), Ecosystem Dynamics in a Polar Desert: The McMurdo Dry Valleys, Antarctica. American Geophysical Union, Washington D.C.

Standard Methods, 1999. 2320 AlkalinityStandard Methods for the Examination of Water and Wastewater. American Public Health Association, American Water Works Association, Water Environmental Federation. Stark, S., O'Grady, B., Burton, H., and Carpenter, P., 2003. Frigidly concentration seawater and the evolutin of Antarctic saline lakes. Australian Journal of Chemistry 56, 181-186.

Stoffyn-Egli, P. and Mackenzie, F., 1984. Mass balance of dissolved lithium in the oceans. Geochim. Cosmochim. Acta 48, 859-872.

Stumm, W. and Morgan, J., 1996. Aquatic Chemistry. John Wiley and Sons, New York, NY.

168

Sverjensky, D., 2006. Prediction of the speciation of alkaline earths adsorbed on mineral surfaces in salt solutions. Geochim. Cosmochim. Acta 70, 2427-2453.

Takamatsu, N., Kato, N., Matsumoto, G., and Torii, T., 1998. The origin of salts in water bodies of the McMurdo Dry Valleys. Ant. Sci. 10, 439-448.

Tanizaki, Y., Shimokawa, T., and Nakamura, M., 1992. Physicochemical speciation of trace elements in river waters by size fractionation. Environ. Sci. Technol. 26, 1433-1444.

Tegt, S., 2002. The Chemical Evolution of Canada Glacier Melt: Supraglacial and Proglacial Waters in Taylor Valley, Antarctica, The Ohio State University.

Tomascak, P., 2004. Developments in the understanding and application of lithium isotopes in the earth and planetary sciences. In: Albarede, F. and Beard, B. Eds.), Geochemistry of non- traditional stable isotopes. Geochemistry of non-traditional stable isotopes, Washington, DC.

Tomascak, P., Hemming, N., and Hemming, S., 2003. The lithium isotopic composition of waters of the Mono Basin, California. Geochim. Cosmochim. Acta 67, 601-611.

Torii, T. and Yamagata, N., 1981. Limnological studies of saline lakes in the dry valleys. Antarctic Research Series 33, 141-159.

Ugolini, F., 1986. Processes and Rates of Weathering in Cold and Polar Desert Environments. In: Colman, S. and Dethier, D. Eds.), Rates of Chemical Weathering of Rocks and Minerals. Academic Press, London.

Viollier, E., Jezequel, D., Michard, G., Pepe, M., Sarazin, G., and Alberic, P., 1995. Geochemical study of a crater lake (Pavin Lake, France) -Trace element behavior in the monimolimnion. Chemical Geology 123, 61-72.

Vitousek, P., Mooney, H., Lubchenco, J., and Melillo, J., 1997. Human domination of Earth's ecosystems. Science 277, 494-499.

VonDamm, K. and Edmond, J. M., 1984. Reverse weathering in the closed-basin lakes of the Ethiopian Rift. American Journal of Science 284, 835-862.

Welch, K. A., Lyons, W. B., Graham, E., Neumann, K., Thomas, J., and Mikesell, D., 1996. Determination of major element chemistry in terrestrial waters from Antarctica by ion chromatography. J. Ion Chromatography A 739, 257-263.

Wilson, A. T., 1964. Evidence from chemical diffusion of a climatic change in the McMurdo dry valleys 1,200 years ago. Nature 201, 176-177.

Witherow, R. A., Lyons, W. B., Bertler, N. A. N., Welch, K. A., Mayewski, P. A., Sneed, S. B., Nylen, T. H., Handley, M. J., and Fountain, A. G., 2006. The aeolian flux of calcium, 169

chloride and nitrate to the McMurdo Dry Valleys landscape: evidence from snow pit analysis. Ant. Sci. 18 497-505.

Wolgemuth, K. and Broecker, W. S., 1970. Barium in seawater. Earth and Planetary Science Letters 8, 372.

Wollast, R., Mackenzie, F., and Bricker, O., 1968. Experimental precipitation and genesis of sepiolite at earth surface conditions. American Mineraologist 53, 645-662.

Wright, A. and Kyle, P., 1990. A. 25. Taylor and Wright Valleys. In: LeMasurier, W. and Thomson, J. Eds.), Volcanoes of the Antarctic Plate and Southern Ocean. American Geophysical Union, Washington, DC.

Yang, Z., Graham, E., and Lyons, W., 2003. Geochemistry of Pyramid Lake sediments: influence of anthropogenic activities and climatic variations within the basin. Environmental Geology 43, 688-697.

170