Factors influencing the dissolution of aragonite in the oceanic water column

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Factors influencing the dissolution of aragonite in the oceanic water column

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Title:
Factors influencing the dissolution of aragonite in the oceanic water column
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Acker, James Gardner
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Tampa, Florida
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University of South Florida
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English
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xxiv, 224 leaves : ill. ; 29 cm

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Aragonite -- Solubility ( lcsh )
Seawater -- Analysis ( lcsh )
Dissertations, Academic -- Marine science -- Doctoral -- USF ( FTS )

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Thesis (Ph. D.)--University of South Florida, 1988. Bibliography: leaves 197-204.

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University of South Florida
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University of South Florida
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024977852 ( ALEPH )
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F51-00172 ( USFLDC DOI )
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FACTORS INFLUENCING THE DISSOLUTION OF ARAGONITE IN THE OCEANIC WATER COLUMN by James Gardner Acker A dissertation submitted in partial fulfillment of the requirements for the degree of Doctor of Philosophy in the Department of the Marine Science in the University of South Florida April, 1988 Major Professors: Dr. Robert H Byrne Dr. Peter R. Betzer

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Graduate Council University of South Florida Tampa, Florida CERTIFICATE OF APPROVAL Ph.D. Dissertation This is to certify that the Ph.D. Dissertation of James Gardner Acker with a major in the Department of Marine Science has been approved by the Examining Commitee on April 14, 1988 as satisfactory for the dissertation requirement for the Ph.D. degree. Examining Commitee: Co-Major Professor: Dr. Betzer Member; Dr. Kent A.

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James G. Acker 1988 All Rights Reserved

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DEDICATION This dissertation is dedicated to my parents, Mr. and Mrs. Robert F. Acker, and to the memory of Dr. Robert M. Garrels whose example of quiet genius will always be remembered. ii

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ACKNOWLEDGEMENTS It is difficult to mention everyone who aided the progress of this research, but several people were vital. Jim Mullins and his co-workers in the Department of Marine Science machine shop constructed elegant solutions to tough problems with minimal instructions. Bob Jolley and Jabe Breland helped cope with problems on research cruises and in the laboratory, and Carolyn Lewis and Renate Bernstein provided moral support, as well as taking over the handling of sediment trap collections. Secretaries Denise Bombard, Carole Cunningham, Ginger Shuert and Lorie Mires were all unsung heroes. Tony Greco provided remarkable photographs with the scanning electron microscope, and Tom Cuba taught me programming skills. The crews and scientific staff of the research vessels Discoverer and Marien-Dufresne also wer e vital to the completion of this work. Dr. Ric hard Feely and his assistant James Gendron of the Pacific Marine Environmental Laboratory were particularly important to the development of this research program. Gillian Rebert-Baldo, S tev e Wolf, and Mike Morris performed important co mplimentary research. Without their contributions, this study would have been more difficult and taken much longer. Most o f all, I wish to recognize Dr. Robert H Byrne and Dr. Peter R Betzer. They pushed me to recognize the merits of double-checking and triple-redundancy, and also encouraged me during moments of doubt and confusion. iii

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LIST OF TABLES LIST OF FIGURES TABLE OF CONTENTS LIST OF SYMBOLS AND ABBREVIATIONS ABSTRACT CHAPTER 1 : INTRODUCTION vii ix xix xxii The Distribution of Caco3 Sediments on the Ocean Floor 7 Aragonite in the Oceanic co2 System 13 The Chemistry of the Oceanic Carbonate System 16 Variation of the Carbonate System in Oceanic Waters 23 Investigations of Caco3 Dissolution in the Ocean 26 Laboratory Investigations of Caco3 Dissolution 31 CHAPTER 2: Caco3 FLUX RESEARCH -METHODS and RESULTS 41 Sediment Trap Experiments -Summer 1982; Summer 1985 45 CHAPTER 3: METHODS OF DISSOLUTION RESEARCH 50 Dissolution Methods -Summer 1982 50 Analysis Algorithm for Potentiometric Data 52 Dissolution Experiments: Spring 1985; Summer 1985; Laboratory 53 Calculations 60 Analysis Algorithm for Spectrophotometric Determinations 65 iv

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CHAPTER 4: RESULTS OF DISSOLUTION EXPERIMENTS Results of Study Results Study Initial of Data Plotted Against (1-n) 2-2-Dissolution Dependence on ([co3 Js -[co3 ]) Evidence of Effects Complication of Functional Dependence Exponential 69 69 79 79 79 85 85 in Long-Term 87 Influence of 6V on Analysis of Dissolution Rate Data 95 Results columnella 95 Range of Exponential Values 103 Dye 103 Analysis 105 of Dissolution Rate Determination 105 of Alkalinity Determinations 107 CHAPTER 5: EFFECT O F SURFACE ALTERATION ON DISSOLUTION RATES 108 n Held Constant at Low and High 108 6s Held Constant at Low and High 109 Effect of in Solutions 111 Physical of Shell 114 Effect of Values of K spa on the Data Analysis 119 Influences of 125 v

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CHAPTER 6 : EVALUATION OF RESULTS 131 Comparison to Previous W ork 131 Data Quality 142 CHAPTER 7: ARAGONITE SOLUBILITY AND THE OCEANIC SYSTEM 146 Investigation of in-situ Dissolution with Laboratory System 146 Use of the Saturation Parameter 6s in Oceanographic Research 152 Processes Affecting the Dissolution of Aragonite in the Ocean 164 Cumulative Factors Affecting Dissolution with Sinking 167 Refined Model of Dissolution with Sinking 168 Estimates of the Aragonite Flux in the Ocean 181 Further Investigations 192 Concluding Remarks 195 LIST OF REFERENCES 197 APPENDICES 205 Appendix I : Calculations of K'spa as a Function of Temperature, Salinity, and Pressure Appendix II: Data from Dissolution Experiments vi 206 219

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LIST OF TABLES TABLE 1 Comparison of aragonite dissolution rate predictions from equations of Morse et. al. (1979), Keir (1980), and Walter and Morse (1985). Dissolution rates are expressed as percent of aragonite mass dissolved per day. [co3 2-Js = 6.469 x 105 mmol/kg sw1 or [co3 2-Js = 7.597 x 105 mmol/kg sw2 TABLE 2 Comparison of dissolution rate equation predictions given in Table 1, compared to the equations generated by potentiometric experiments in Byrne et. al. (1984). These are Equations 37 and 38 in the text of the dissertation. Dissolution rates are expressed as percent of shell mass dissolved per [co3 2-Js 6.469 x 105 mol/kg sw1 [co3 2-Js = 7.597 x 105 mol/kg sw2 or [co3 2-Js 38 8.754x 105 mol/kg SW3 72 TABLE 3 Values of n and residual sum-of-squares (determined by least-squares analysis) for dissolution rate vs. 6.s datasets, AND n and residual sum-of-squares for dissolution rate vs. (1-n) datasets. 6.V = -31.3 cm3 /mol. (Subsequent analyses employed a different value of 6.V, the molar volume change for aragonite dissolution.) vii 86

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TABLE 4 Values of n and residual sum-of-squares (determined by least-squares analysis) for dissolution rate vs. datasets, AND n and residual sum-of-squares for dissolution rate vs. (1-G) datasets. = -39.5 cm3 /mol. TABLE 5 Residual sum-of-squares values for non-linear least-squares curve fits for rate vs. datasets 1-15 (shown in tables 3 and 4). Data was generated for three 101 values of -31.3, -37.0, and -39.5 cm3 /mol. 102 TABLE 6 Data on pteropod shell settling rates. Each experiment will be described separately, as several different methods were employed. 170 TABLE 7: Values of K' and K' at 25 degrees spc spa Centigrade and 35/00 salinity. The studies which determined these values are also shown. The units for all values are (mol/kg) 2 208 TABLE 8: K' values obtained between 2 and 5C, at spc 35/00 and 1 atmosphere pressure. TABLE 9: Values of for aragonite and calcite dissolution, all at 2C. The studies from which these values were obtained are also shown. Units are cm3/mol. viii 210 217

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LIST OF FIGURES Figure 1 (a-f) Representative examples of pteropod and heteropod shells. a. Cuvierina columnella b. Cavolinia tridentata c. Limacina bulimoides d. Limacina helicina e. . Clio pyramidata (with juvenile helicina shown to scale) f. heteropod Atlanta megalope. Heteropod shell morphology is quite similar within this genus. (b-f provided by John McGowan of the Scripps Institute of Oceanography.) Figure 2 a (from Broecker and Peng, 1981) Distribution of deep-sea calcite sediments ("foraminiferal ooze") on the ocean floor. b. (from Berner, 1977) Map of . the distribution of pteropod ooze in the oceans Figure 3 (from Berger, 1975) Profile of dissolution rate of optical calcite on a moored buoy in the central Pacific demonstrating the existence of a level of rapid increase of dissolution near 3 ,700 meters. Results of Peterson (1966) Figure 4 (from Berger, 1975) Profile of the preservation aspect ("solution index") of foraminifera in the central Atlantic, demonstrating an abrupt transition from well-preserved to poorly-preserved foram assemblages (foram "lysocline"). Points plotted are averages (mean N = 7). Results of Berger (1968). ix 3 9 27 29

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Figure 5 (from Morse and Berner, 1974) Schematic plot of rate of dissolution versus 6pH. The inset is an enlargement of the region near equilibrium demonstrating the presence of a pronounced discontinuity which is otherwise lost on the scale of the main diagram. This discontinuity is believed to represent the chemical lysocline. Figure 6 (from Morse et. al. 1979) The log of the . dissolution rate versus the log of ( 1-0) for synthetic aragonite and pteropods. Note the single point for pteropod shells, and the substantially reduced dissolution rate of the shells. Figure 7 -Cruise tracks and stations for North Pacific cruises on board the R/V Discoverer. a. May-June 1982. b. June-July 1985. Figure 8 Cruise track of the R/V Marien-Dufresne in the southern Indian Ocean and circumpolar Antarctic Ocean (INDIGO I cruise), February-March 1985. Figure 9 (from Betzer et. al. 1984a) Sediment trap of the design used by Betzer et. al. (1984b) for North Pacific . cruises on-board the R/V Discoverer. Figure 10 (from Betzer et. al. 1984b) Depth . distributions of pteropod fluxes (in milligrams per square meter per day) from the western North Pacific. Aragonite saturation state is shown for 100 percent saturation( __ ), 80 percent saturation ( ----), and 60 percent saturation (--) Calculated fluxes do not include small 10 percent) contributions from pteropod fragments. X 34 36 43 44 46 48

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Figure 11 of and pathlength cell (shown with ) used in this study. 12 effects of solution on dissolution of Note that the 60-85 minutes of is equivalent to initial dissolution 1 3 -Schematic of analysis technique. Initial data was collected by with Apple lie samples were collected by bottle cast seawater samples in analyzed by titration. Data experiments was and filed into and analyzed by custom Statistical analysis of data was by mainframe and Statistical Analysis S ystem 14 Byrne et. al., 1984) The solubility obtained in investigations as a function of equilibration time. The p = K' /K' a direct o f K' the spa spc spc solubility of calcite obtained by Ingle et. al. (1973) (35 /00 salinity, 25C) and K'spa values (35 00 salinity, 25C) obtained by: e, Plath (1979); 6, (1976); o, Macintyre (1965); 0, et. al. (1980). K's o . pc values obtained in at 25 C and 35 00 salinity to within 5%. xi 54 57 68 71

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Figure 15 (from Byrne et. al., 1984) The cumulative % dissolution of pteropods settling at a mid-range rate, 1.4 em sec-1 at two of our station locations in the North Pacific Ocean. Significant dissolution begins at shallow depths at the northernmost station (50N, 167E). At 35N, 165E, dissolution becomes appreciable below 1000 m. As the influence of dissolution on settling rate is not considered, the cumulative dissolution shown here should be considered as a lower-bound estimate. Figure 16 (from Byrne et. al. 1984) The predicted c umulative % dissolution of pteropod shells in the North Pacific versus depth for representative initial settling velocities. The initial settling rates (in centimeters per second) appropriate to each species are based on results obtained in our laboratory. The scales adjacent to each pteropod species represent a length of 1 mm. b. A single pteropod species ( Limacina helicina) was observed in the highly-undersaturated waters at the northernmost station. Settling rates were determined by several methods (see table 6). Observation of the shells throughout their descent indicated a uniform settling rate apparently unaffected by turbulence. Figure 17 -Dissolution rate vs. aragonite saturation, here expressed as Q .t Experimental sequence used a aragon1 e single _f. tridentata shell, mass 45.608 mg. Performed on-board the R/V Marien-Dufresne, March 1985. This figure shows high-and low-pressure regions (shaded areas, with pressures in psi) in the same (1-n) range. xii 75 77 80

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Figure 18 Data from figure 17 shown plotted against 22-[C03 JsCco3 J (=lis). "L" designates data points from the last two experiments in the sequence. Figure 19 6(1-n)/6P and 6(6s)/6P, as calculated for a hypothetical sample of seawater (35 /00 salinity, 5C, TA 2 . 400, EC02 = 2.3387), which is exactly saturated at atmospheric pressure. 6(6s)/6P increases with increasing pressure, in contrast to 6(1-n)/6P, which decreases with increasing pressure. Figure 20 Data from the first three days of experimentation of a 9-day, 10-experiment sequence with one assemblage of a C. tridentata and C. gibbosa shell. Figure 21 Data from the next two days (three expe riments) of the 9-day sequence using the same shell assemblage. Figure 22 Data from the final four days of 9-day sequence using the same shell assemblage. Figure 23 Increasing rate curves from 9-day, 1 a -experiment sequence. Curves were generated by least-squares analysis of data from Figures 20 (1), 21 (2), and 22 (3). xiii 82 84 88 89 90 91

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Figure 24 -Increasing rate curves from second 9-day, 9-experiment sequence. , first three days; +, second three days; o, final 3 days. This sequence also shows increasing dissolution rates with extended exposure to undersaturated seawater. Figure 25 Variable pressure aragonite dissolution rate data from Cuvierina columnella experiments is examined using Eqns. (17) and (35), and Eqns. (36) and (35). For each model, the residual sum of squares for our fitted data is shown as a function of 11V. Figure 26 -Dissolution rate data fr-om C. columnella 2-2 -experiments shown as a function of [co3 Js -[co3 ] ( /1s), using our best-fit /1V estimate of -37.0 cm3 /mole. Figure 27 Results from experiments in which the phenol red concentration was varied between 1/2 and 2x the nonnal experimental concentration. 1/2 -one-half normal concentration; N-normal concentration (8 x 10-6M); 2xtwice normal concentration. Figure 28 Exper-imental sequence during which shells were stor-ed in highly-undersatur-ated seawater solution between exper-iments showing strong effects of advancing dissolution. o, Day 1; 0, Day 2; +, Day 3; X, Day 4; o, Day 92 97 98 106 5. 112 Figure 29 (from Morse, 1986) Effect of decreasing grain size on the the bulk solid solubility of calcite. xiv 115

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Figure 30 (top) Limacina inflata shell taken from 2170 meters in the North Pacific Ocean, showing the effects of dissolution on the surface and the roughened nature of the partially-dissolved surface. (middle) Ridges on a Cavolinia tridentata shell used in dissolution experiments, demonstrating initial dissolution effects. ( bottom ) Closeup of ridge surfaces, showing roughened surface characteristic of exposed aragonite rods. Figure 31 -a. Data from experiments using C. tridentata shell on-board the R/V Marien-Dufresne, from first days of + data from the last two days of experimentation. b. Data from the last two days adjusted with increased value of K'spa' (This data also appears in Figure 1 8 ) Figure 32 Data from initial experiments using a C tridentata shell, mass 12.502 mg. (+) Last two days of b. Data from last two days adjusted with increased value of Figure 33 -a. Data from 9-day, 10-experiment sequence divided into three groups corresponding to Figures 20, 21, and 2 2. 0 first three days; +, next two days; +, last 4 days. b Data adjusted with increased values of K'spa XV 117 120 122 123

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Figure 34 -a. Idealized experiment in which aragonite d issolution rates are determined instantaneously over a continuous under-saturation range, showing a theoretical linear relationship between dissolution rate and D.s. b. Idealized experiment in which the dissolution kinetics of aragonitic material previously exposed to undersaturat ed seawater are determined instantaneously. Idealized experiment illustrating t he effects of rate determination at discrete levels of under-saturation, causing enhanced dissolution rates due to the increased surface area of the partially-dissolved surface. 127 Figure 35 (from Morse et. al., 1979) The ratio of measured rate of dissolution per gram to the initial dissolution rate as a function of extent of dissolution for pteropods and calcitic Pacific Ocean sediment ( Morse, 1 978). 133 Figure 3 6 -Lei Chou's plot of dissolution rate vs. pH, shown with curves from Plummer et. al. (1978). T h e square shows Chou's data in the near-saturation region which are comparable to the results in this dissertation. Figure 37-Chou's data plotted against (1-n), t a k en from the square in Figure 36. Figure 38 -Dissolution rate and n pr-ofile at 50N, 145W in the North Pacific Ocean. Each dissolution rate was obtained using a single f tridentata shell, mass 16.2 .1 . mg. The calculated n profile cor-r-esponds to the satur-ation state of our shipboard experiments which were per-formed at 5C. Pressures cor-r-esponded to t h e depths fr-om w hich the 139 140 seawater samples were obtained. 148 xvi

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Figure 39 Plot of measured dissolution rates at four stations in the North Pacific Ocean, July 1985. Experiments were performed at near in-situ pressures and 5C. The in-situ temperature range was 6.75 to 4.1C. Measured rates coincide with the calculated saturation horizons, increasing rapidly below the 100% saturation horizon. 151 Figure 40 -a. Saturation state vs. depth profile for Station 3 (July 1985), 28N, 155W in the North Pacific Ocean. The 6s axis has been shifted so that the profiles approximately coincide in the upper water column, to show the divergence of the two parameters with increasing depth. b. Saturation state vs. depth profile for Station 9 (July 1985), 42N, 148W in the North Pacific Ocean. The has been shifted in the same manner as for Figure 40a. 155 Figure 41 -Plot of in the E. North Pacific Ocean, July 1985. (Unpublished data provided by J Gendron) 156 Figure 42 -Plot of n in the E North Pacific Ocean, July 1985. (Unpublished data provided by J. Gendron) 157 Figure 43 -Plo t of in the W. North Pacific Ocean, June 1982. (Data from Feely et. al., 1984) 158 Figure 44 Plot of n in the W. North Pacific Ocean, June 1982. (Data from Feely et. al., 1984) 159 xvii

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Figure 45 Plot of 6s in the S. Atlantic Ocean. (GEOSECS data) Figure 46 Plot of G in the S. Atlantic Ocean. (GEOSECS data) Figure 47 Plot of 6s in the S. Indian Ocean, February-March 1985. (INDIGO I data) Figure 48 Plot of G in the S. Indian Ocean, February-March 1985. (INDIGO I data) Figure 49 -Model of dissolution with sinking, using data from 50N, 167W in the North Pacific Ocean. Dissolution rate equations used to generate curves : I. Rate== II. Rate== .004 III. Rate == .01 (6s) 1 IV .. Rate = .oi (6s) 1 with 6V for aragonite dissolution= -37.0 6 V = -31.3 cm3 /mol for the first three curves. Figure 50 -Model results for various settling rates from 1 to 2 em/sec Equation: Rate== 01(6s) 1 Figure 51 -Model results for settling rate of 1 .5 em/sec. Equation : Rate = Data from Northern 0 Pacific stations 6 9, and 13, (July 1985) and 50S in the Indian Ocean. xviii 160 161 162 163 174 179 180

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A a A min A max atm BA B(OH)3 B(OH)4 BT CA Ca2+ Caco3 Cl C02 D E H+ H202 HCl HgC12 Hg Hg(II) HgC 1 4 LIST OF SYMBOLS AND ABBREVIATIONS absorbance activity minimum absorbance maximum absorbance atmospheres borate alkalinity boric acid borate ion total boron concentration Cat'bonate alkalinity calcium ion calcium carbonate chlorine I chloride ion centimeter cubic centimeter carbonate ion critical carbonate ion concentr"ation carbonate ion concentration at saturation carbo n dioxide total amount of dissolution east hydrogen ion hydrogen peroxide hydrochloric acid mercuric chloride mercury mercuric ion common ionic association of mercuric and chloride ions in seawater xix

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K I 1 K I 2 K I B Klsp Kl spa K1spc kg 1 ln log M m 2 m m3 Mg mg mm ml mmol mol N NBS nm NOAA OH p watercarbonic acid bicarbonate ion rate constant apparent dissociation constant of carbonic acid apparent dissociation constant of bicarbonate ion apparent dissociation constant of boric acid apparent solubility product apparent solubility product of aragonite apparent solubility product of calcite kilogram liter natur-al logarithm base 10 logarithm molar concentration molal concentration square meter cubic meter magnesium milligram millimeter milliliter millirrole mole north National Bureau of Standards nanometers National Oceanic and Atmospheric Administration hydroxide ion pressure XX

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pH pK1 psi R R/V RSS s SAS SEM sw T TA T.A.A.F TRIS w yr Miscellaneous symbols 6G 6K t.s 6V K lleq llM \liTl llffiOl EC02 n o;oo negative logarithm of the hydrogen ion concentration negative logarithm of the dissociation constant pounds per square inch gas constant or dissolution rate research vessel residual sum-of-squares salinity I surface area/ south Statistical Analysis System scanning electron microscope seawater temperature total alkalinity Territoires Australes et Antarctiques Francais Tris(hydroxymethyl)aminomethane west year free e nergy molar compressibility change 22([C03 ]s [C03 ]) molar volume change for a reaction alternative rate constant microequivalents micromolar (10-6 molar) mic ron (10-6 meter) micromoles total inorganic carbon content 2+ 2[Ca ][co3 ]/K's May be defined P calcite or aragonite using appropriate degrees Centigrade parts per thousand xxi

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FACTORS INFLUENCING THE DISSOLUTION OF ARAGONITE IN THE OCEANIC WATER COLUMN by James Gardner Acker An Abstract of a dissertation submitted in partial fulfillment of the requirements for the degree of Doctor of Philosophy in the Department of Marine Science in the University of South Florida April, 1988 Major Professors: Dr. Robert H Byrne Peter R. Betzer xxii

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The dissolution kinetics of biogenic aragonite in seawater were investigated by experimentally simulating oceanic temperature and pressure conditions. The equation Rate = k xn relates aragonite dissolution rates to the degree of seawater undersaturation, expressed by x. Previous research used x = (1-0), where n is the ratio of the product of calcium and carbonate ion concentrations in seawater to the apparent solubility product of aragonite in seawater My analyses demonstrate that aragonite dissolution rates are described more 22-successfully with x = ([co3 Js-[co3 ]), the difference between the carbonate ion saturation concentration and the in-situ carbonate ion concentration. Experiments also indicated progressively increasing dissolution rates when the same shells were used in consecutive rate determinations. The dissolution rate increase apparently results from shell surface alteration. Partially-dissolved shells taken from sediment traps deployed in undersaturated waters of the Pacific Ocean showed microscopic alterations resulting from dissolution. My observations suggest that in the absence of surface alteration, the dissolution rate equation would be linear in the parameter x = ([co32-JsHowever, natural surface alterations cause the rate equation appropriate to descriptions of aragonite dissolution in the ocean to be non-linear in x. Oceanic cross-sections of ([co32-Js -[co32-]), and a new model of the aragonite dissolution process, allow prediction of the fate of sinking pteropod shells in the ocean. For the northern Pacific Ocean, xxiii

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the model predicts that shells with slow rates of descent will be completely dissolved before reaching the ocean floor. Due to the high degree of undersaturation in these waters, aragonite may constitute a source of alkalinity in the northern Pacific water column. In previous studies, aragonite fluxes may have been underestimated due to selective dissolution of aragonite in deep sediment trap deployments. In contrast, pteropod behavior may lead to overestimation of aragonite fluxes in shallow traps. However, estimates of the aragonite flux based on data from shallow trap deployments accord with independent models of the oceanic co2 system. The operation of biological processes (such as predation) may act to increase the delivery rate of biogenic aragonite to deep oceanic waters. Abstract approved: 1 'X .. I ..._.Y .. Co-major Professou --1 _,_ ___ ..__,_ ,._ ..... t 1 > I o I V Date f Approv 1 xxiv

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CHAPTER 1 : INTRODUCTION "The sea never changes and its works, for all the talk of men, are wrapped in mystery." Joseph Conrad, Typhoon In the oceanic system, the formation of calcium carbonate (Caco3 ) involves the biology of numerous organisms, ranging from massive coral reefs to microscopic phytoplankton. Several different taxons form shells or tests made of Caco3 and sedimentary deposits composed of these shells are found over wide areas of the ocean floor. The ultimate fate of caco3 in the oceans is strongly coupled to the chemistry of seawater. In some areas, Caco3 accumulates in large quantities on the ocean floor. In other areas, Caco3 dissolves so rapidly that calcareous deposits cannot accumulate. The laws that govern the stability of Caco3 minerals in the oceans have been the subject of research for many years, and will continue to be important in the future as mankind's impact on the oceans increases. The burning of fossil fuels for energy production by mankind has added large quantities of anthropogenic co 2 to the atmosphere. Some of this "fossil fuel C02 is absorbed by the oceans, where it will interact with caco3 already present in the oceans. An additional facet of the eaco3 story in the ocean derives from the mineral's two polymorphic forms, calcite and aragonite. Although they have the same chemical composition, the differing mineralogy of the two minerals (calcite has a rhombohedral structure, while aragonite is orthorhombic) endows them with different chemical

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properties. One notable difference is aragonite's greater solubility (Macintyre, 1965). This simple factor has large-scal e consequences, as oceanic sediments composed of calcite are widespread, but sediments containing a high percentage of aragonite are small in extent and relatively rare. Although variations in biological productivity also affect the distribution of sediments, the greater solubility of aragonite is the primary factor controlling the appearance or absence of aragonite on the ocean floor (Berner, 1977). Marine organisms may form Caco3 hard parts or shells composed of either calcite or aragonite. In the shallow marine environment, corals and molluscan shells are composed primarily of aragonite, and ca l careous algae can form both polymorphs, depending on species. (MacKenzie et al., 1983). Calcareous algae may manufacture calcite with a high mole percentage of magnesium (Mg) which is referred to as "high-Mg" calcite (McKenzie et al., 1983). In the pelagic ocean, five c lasses of organisms manufacture Caco3 Coccolithophorids, foraminiferans, and ostracods form "low-Mg" calcite shells, which will be referred to simply as ca lcite. Ostracods are only lightly calcified, and rarely contribute calcite to deep-sea sediments, probably due to rapid dissolution of the fragile shells once they reach the sediment-water interface. The other classes of planktonic organisms which form a eaco3 shell are the pteropods and heteropods. In contrast to the other three classes, their shells are made of aragonite. Figure 1 shows several examples of pteropod shells, demonstrating the variation in the size and shape of the shells. Data on the oceanic distributions of pteropods and heteropods is still limited. The organisms are known 2

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1 (a-f) examples of and shells. a colurnnella b. Cavolinia c. Limacina bulimoides d. Limacina helicina e. Clio pyramidata (with helicina shown to scale) f. heteropod Atlanta megalope. shell morphology is quite similar within this genus. (b-f by John McGowan of the Institute of

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1111111 a Cuvierina columnella Limocino helicina A Limacino helicina B Allonlrl lmm ....._... c d 4 b I \ I Cavolima tndentata L i moctna bulimoides e .

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to exhibit seasonal changes in abundance, and pteropods are generally more numerous than heteropods. In some regions, pteropods are a minor food source for whales (Be' and Gilmer, 1977). However, heteropods can outnumber pteropods in some oceanic regions (V. Fabry, personal communicat i on) Several factors will determine the ultimate fate of the Caco3 shell which is formed by a planktonic organism. Coccolithophores are consumed in large amounts by zooplankton, primarily copepods, which excrete the caco3 forms packaged in fecal pellets. The resistant covering of the pellet, combined with its rapid sinking rate, promotes transport of the Caco3 contents to the sea floor while protecting the forms from dissolution (Honjo 1976). In highly productive areas, large amounts of Caco3 may be contributed to the sediments in this manner (Honjo and Roman, 1978; Honjo et al., 1982). Entrainment in the amorphous organic materials known as "marine snow" may also be a transport mechanism for coccoliths and coccospheres (Silver and Alldredge, 1981). Predatio n does not have as great an impact on the fate of foraminifera and pteropod shells as for coccoliths. However, pteropods are so numerous in some areas that they are a food source for whales and fish (Be' and Gilmer, 1981). Although it is only an isolated example, a large quantity of Cavolinia longirostris shells were found in the gut of a triggerfish from the Gulf of Mexico (S. Condileri, pers. communicat i on) The fate of the shells consumed by predators is unclear. Mortality of foraminifera and pteropods is the primary reason that the shells will begin to sink in the water column. Observations 5

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6 at 50N, 165W in the North Pacific Ocean (Station "Papa") during July 1985 indicated that juvenile Clio pyramidata and juvenile Limacina helicina were prevalent. However, approximately one year later, adult Limacinas with larger shells were prevalent at the same location (V. Fabry, unpublished data). Therefore, the sizes of shells delivered to the water column can vary substantially ( 63 J.Jffi to 3 em, Be' and Gilmer, 1977). Once the sinking process has started, the factors that will influence the fate of the shell are strictly physical and chemical in nature. The mass, volume, and shape of the shell will be the main factors which determine the sinking rate. Seawater viscosity, which is influenced by temperature and pressure, will also affect the sinking rate. Initial dissolution of the shell may be caused by bacterial oxidation of organic matter within the shell. This oxidation will make the micro-environment within the shell more acidic, and could lead to some dissolution of the shell's interior. (This process is implied from the observations of Gardner et al., 1983). Subsequently, as the shell sinks, it will be exposed to conditions in the deep water column which will induce dissolution of Caco3 ("corrosive" conditions). Honjo (1977) states that the sinking rate of isolated coccoliths is so slow that the plates would be rapidly dissolved under conditions Hhich favor eaco3 dissolution. However, the primary mode of transport for coccoliths (in fecal pellets) protects them from dissolution. The more rapid sinking rate of foraminiferal tests allows them to reach the ocean floor virtually unaffected by dissolution, according to

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Honjo. In most oceanic regions, aragonite will also be relatively unaffected by dissolution during settling. Hm.,rever, in the North Pacific, dissolution of aragonite may be significant at shallow depths. In this region, dissolution can rapidly decrease the mass of an aragonitic shell, which will reduce the sinking rate, and subseque ntly increase the exposure time to the corrosive waters. This process may lead to the destruction of aragonite mass in the Pacific water column (Byrne et al. 1984) Further discussion of th.is important process will be found in this dissertation. Once on the bottom, the shell or test will be subject to further corrosive attack, which may be mediated by the presence of the "diffusive sub layer" (Schink and Guinasso, 1977; Boudreau and Guinasso, 1982) In this layer, usually one millimeter or less in thickness, molecular diffusion is the primary mode of mass transport. The reduced circulation in this layer retards the transport of dissolved ions from the shell surface and into the overlying waters, where advective transport becomes important. A further impediment to dissolution may occur in areas with high rates of sedimentation, where rapid burial rates can decrease the residence time of the shells at the sediment-water interface (Berner, 1970) The Distribution of Caco2 Sediments on the Ocean F loor A major uncertainty concerning the distribution of calcite and aragonite on the sea floor is whether their occurrence in sediments is proportional to the flux of the minerals in the water column, or 7

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whether the sedimentary aragonite/calcite ratio has been altered by dissolution. Deep-sea sediments consisting mainly of calcite (foraminiferal ooze) cover wide areas of the ocean floor (Figure 2a), whereas pteropod ooze, which is composed primarily of aragonite, is only found in isolated regions (Figure 2b) Shallow aragonitic sediments resulting from coralline and algal aragonite production may be found in near-shore areas. On oceanic rises, a progression of aragonitic (pteropod) ooze to calcitic (foraminiferal) ooze to red clays is usually found with increasing depth (Murray and Hjort, 1912, cited in Berger, 1975). The primary reason for the variation in deep-sea eaco3 sediments is believed to be the greater solubility of aragonite relative to calcite (Berner, 1977). However, the different trophic levels of coccolithophores, foraminifera, and pteropods, as well as overall rates of surface productivity, will influence the distribution of the sediments. Severa l "horizons" have been described that indicate factors which influence the distribution of calcite and aragonite in sediments. The shallowest horizon that will be encountered is the saturation horizon, where the water is exactly saturated with respect to calcite or aragonite. Because of aragonite's greater solubility, the aragonite saturation horizon lies above the calcite saturation horizon. Proceeding deeper, the next horizon that will be encountered is the lysocline, which is the depth at which the effects of dissolution on carbonate forms become apparent. The experiment of Peterson (1966) demonstrated a marked increase in the weight loss due to dissolution 8

PAGE 35

Figure 2 -a. (from Broecker and Peng, 1981) Distribution of deep-sea calcite sediments ("foraminiferal ooze") on the ocean floor. b. (from Berner 1977) Map of the distribution of pteropod ooze in the oceans. 9

PAGE 36

of suspended calcite spheres at a certain depth. This depth was originally termed the "lysocline". Berger ( 1975) states "the term 'lysocline' was originally coined in the belief that the striking change in foraminiferal composition at the interface of the Antarctic Bottom W ater and the North Atlantic Deep Water constitutes a trace of acceler-ated dissolution in the water column; that is, a Peterson l evel." 10 In subseq uent work, Berger (1968; 1975) defined the foraminiferal lysocline, which is the depth at which the maximum change in the ratio between the "easily dissolved" and the "resistant" foraminifera occurs. It appears that a large degree of calcite dissolution has taken place, perhaps 80%, in the foraminiferal lysocline zone. Berger's horizon has also been termed the sedimentary lysocline. Morse (1974b) defined the "chemical lysocline" as "the depth corresponding to the under-saturation at which a distinct increase occurs in the rate of dissolution of calcite with increasing undersaturation." The "original" lysocline, based on Peterson's data, is similar to Morse's term. Because the c hemical lysocline corresponds to a certain level of undersatur-ation, this horizon may be defined according to temperature, pressure, and seawater composition. This definition is in contrast to the sedimentary lysocline, which is defined on the basis of organism assemblages in sediments. The third important horizon is the carbonate compensation depth (Bramlette, 1961), or CCD. Two compensation depths may be defined; the calcite compensation depth and the aragonite compensation depth. In either case, the horizon is defined as the point at which the delivery rate of the mineral to the ocean floor is exactly balanced by

PAGE 37

11 the r-ate of miner-al dissolution at that depth. Inher-ent in this definition is the assumption that the sedimentation r-ate of non-carbonate constituents is essentially constant (Broecker and Peng, 1982) Below the compensation depth, significant accumulation of the miner-al wil l not occur because the rate of loss d u e to dissolution exceeds the delivery rate. Due to aragonite' s greater solubility, the aragonite compensation depth lies much shallower than the calcite compensation depth. The shallowness of the aragonite compensation depth means that sediments containing a large percentage of aragonite will only be found on shallow regions of t he sea floor. By examining the Holocene sedimentary record, Broecker and Takahashi (1977) correlated dissolution in carbonate sediments with the chemical state of the over l ying seawater. The authors state that no measurable calcite dissolution will occur until the seawater carbonate ion content drops below the "critical carbonate ion content", given by this equation: ( 1 ) 93 e 0 14(z-4) z is depth in meters /100. Broecker and Takahashi then used the shape of the sedimentary "transi t ion z o ne" profile (where the caco 3 content of the sediment decreases rapidly with depth) to model the dissolution rate of the sediments These rates were modeled as a function of the critical c arbonate ion content, with the equation (2) Rate 2 rnoles/m /yr

PAGE 38

f is the fraction of calcite in the bottom sediments. (A note added by the editor to Broecker and Takahashi (1977) indicates that the value of the coefficient should be 0.0125 rather than 0.025 ) 12 The map shown in figure 2 (Berner, 1977) shows the areas in which pteropod ooze (defined as Caco3 sediments in which the majority of the Caco3 is in the form of pteropod shells) are found on the ocean floor. It is important to note that aragonite sediments are virtually absent from the entire Pacific Ocean floor, and essentially all of the circumpolar Antarctic Ocean as well. In a related paper (Berner et al., 1976), the progressive c hange in carbonate sediments with depth on the Bermuda Pedestal was evaluated. Relatively shallow sediments on the Pedestal are predominantly aragonite. Calcite becomes the dominant mineral as depth increases, and the surviving aragonitic debris shows increasingly severe effects from dissolution with increasing depth. The difference in the saturation state of the deep Atlantic and Pacific Ocean waters with respect to calcite is believed to affect the occurrence of calcite sediments (primarily foraminiferal oozes) on the floor of each ocean. Only a few areas of the Atlantic are deep enough to lie below the calcite compensation depth, and therefore carbonate sediments containing at least 25% calcite are common in the Atlantic. Along the crest of the Mid-Atlantic Ridge, sediments consisting predominantly of calcite (greater than 80% calcite, Biscaye et al., 1976) are common. In the Pacific, the area of the ocean floor covered by calcite sediments is small relative to the Atlantic. Predominantly calcitic

PAGE 39

13 sediments are found along the crest of the East Pacific Rise and in the southwestern Pacific, but there are virtually no calcite sediments found on the abyssal areas of the northern Pacific Ocean. In general, substantial dissolution effects on s e diments are evident at 5000 meters in the Atlantic and at least 1500 meters shallower in the Pacific. (Berger, 1975). Aragonite in the Oceanic co2 System Due to its relatively minor occurrence in recent marine sediments, aragonite has not been considered of major importance in the global geoche mical system. In contrast, calcite has long been known to be important in the global and oceanic carbon dioxide (C02 ) system Aragonite's greater solubility causes the volume of the ocean that is undersaturated with respect to aragonite to be larger than the volume \vhich is undersaturated with respe c t to calcite. Furthermore, aragonite' s greater solubility means that oceanic waters which are undersaturated with respect to calcite have a higher degree of undersaturation relative to aragonite. These two factors would appear to indicate that aragonite dissolution might be a more i mportant factor in the chemistry of the water column than calcite dissolution. The more labile nature of aragonite relative to calcite may be offset by the fact that the forms of aragonite that are produced in the pelagic ocean, pteropod shells, are much larger than calcitic forms Aragonitic shells 1 mm to 1 em) are generally larger in size than foraminiferal tests 100 to 1 mm) and are much larger

PAGE 40

than coccoliths (1 to 10 The size difference, combined with the slightly greater density of aragonite (2.94 grams/cubi c centimeter) than calcite (2 7 1 grams/cubic centimeter) indicates that the settling rates of aragonitic shells will exceed the settling rates of foraminiferal tests. The settling rates of aragonitic shell s will be orders of magnitude greater than the settling rates of isolated coccoliths. However, small pteropods settle at rates comparable to foraminiferal tests. Adelseck and Berger (1975) state that the rap i d settling rates of large pteropod shells allowed their collection at great depth in water that is substantially undersaturated with respect to aragonite. Therefore, in spite of the fact that large volumes of the ocean are undersaturated with respect to aragonite, the rapid downward transport of pteropod shells ap pears to indicate that the majority of a rag onite dissolution in the oce an will take place on the sea floor. Other factors which affect the dissolution of biogenic caco3 are the surface area and porosity of the various forms Since coccoliths are m uch smaller than pteropod shells, their active surface area is much greater. Furthermore, porous foraminiferal tests will have a greater surface area than an equivalent mass of pristine, non-porou s pteropod shells. The a bundan ce of pteropods relative to foraminifera and coccolithophores is undetermined As a percentage o f recent marine carbonate sediments aragonite is a minor constituent. Many sediment trap collections have also shown a markedly greater flux of calcite to the deep ocean than aragonite. In contrast, net collections of plankton indicate that pteropod shells are "caught" in approximately 14

PAGE 41

15 the same numbers as foraminifera (Adelseck and Berger, 1975; Berner, 1977). Also, in net collections the mass of aragonite usually exceeds the calcite mass (Berger, 1978). Furthermore, sediment pteropod /foraminifera ratios (by numbers of each class) cited in Berner (1977) are approximately 1:1 in sediments collected from above the aragonite saturation horizon. However, much shorter life-spans for foraminifera than pteropods (Be' and Gilmer, 1976; Kobayashi, 1976; Harbison and Gilmer, 1987) and the grazing of zooplankton on coccolithophores (Honjo, 197 6) indicate that the turnover rate for foraminifera and coccolithophores should be much faster than for pteropods. Absolute turnover rates have not been determined for pteropods, but Berger (cited in Berner, 1977) indicates that their turnover rates are slower than the turnover rates for calcite-producing organisms. Higher turnover rates should result in greater numbers of sinking calcitic particles delivered to the deep ocean, compared to the numbers of sinking aragonitic particles. Note that the observation of a 1 :1 pteropod /foraminifera ratio in sediments which are unaffected by dissolution conflicts with this expected greater rate of calcite delivery. Thus, aragonite and calcite may be present in roughly equal amounts (by mass) as living plankton in the ocean (Berger, 1978). Pteropods and foraminifera are also found with roughly equal numerical abundances in sha llow marine sediments (Berner, 1977) Yet aragonite appears to disappear rapidly in the deep oceanic environment, while calcite survives to form widespread sediment deposits. If dissolution is the eventual fate of nearly all of the aragonite produced in the pelagic surface ocean, the factors governing the dissolution process

PAGE 42

16 need to be well understood. The Chemistry of the Oceanic Carbonate System In the above discussion, reference is made to the "saturation state of the oceans with respect to calcium carbonate". Also, sane regions of the ocean are referred to as "undersaturated". Fundamental chemical relationships which define the oceanic carbonate system support these statements. These chemical relationships must be defined prior to a discussion of calcium carbonate dissolution kinetics. In the simple case of solid calcium carbonate dissolving in water, the mineral phase will be dissociated into its constituent ions. This reaction can be represented as (3) 2+ 2Ca (aq) and co3 (aq) At equilibrium, the rate of the dissolution reaction, shown above, and the reverse reaction (precipitation) \-lill be the sa.11e. The equilibrium constant K for dissolution is defined in terms of the activity of each species: ( 4) K aeaco 3 The activity of a pure mineral phase is equal to one Therefore,

PAGE 43

17 the expression above becomes the definition o f the solubility product, K for eaco3 : sp ( 5) In pure water, the activities of the ions would be equivalent to their actual concentrations. At tne theoretically low concentrations of the ions in pure water, there are no constituents which can interact with the ions, so the concentration of the "free" ions is the same as the total ion concentration. In seawater, however, several interactions due to the presence of several dissolved ionic species reduce the ionic activities relative to the actual ion concentrations, and also cause only a percentage of the ions to be in the "free" state. It has become customary for the oceanographic community to use 'apparent' or stoichiometric solubility products, K sp in reference to mineral solubility in seawater. K' is defined as the product of the total sp concentrations of each ion at saturation: (6) K sp K' may be defined in terms of several concentration units. The most sp common units used in oceanographic research are moles per kilogram of seawater (mol/kg sw), but either molarities (moles/liter) or molalities (moles/kg H 2 0) may also be used. A small "m" with an ion subscript (such as mca2+) indicates total molal concentration in this work, and brackets indicate total concentrations in moles per kilogram of seawater.

PAGE 44

The seawater saturation state with respect to Caco3 has been expressed in terms of the parameter n (Edmond and Gieskes, 1970; Morse, 1978), which is defined by the solubility product and the ion concentration product: (7) K' sp Note that when the seawater is exactly saturated with respect to Caco3 the value of n will be equal to 1. The quantity will be greater than 1 in the case of supersaturation and less than 1 in the case of undersaturation. A closely-related quantity that is particularly useful to indicate the degree of seawater undersaturation is (1-n) The value of the q uantity (1-SI) increases as the degree of seawater undersaturation increases. 2-2-The quantity !J.s = [ co3 ]s -[co3 ] provides an alternative direct measure of seawat e r saturation state ( Takahashi et al, 1980; Acker et al., 1987). This quantity is the difference between the carbonate ion concentration at saturation and the actual (or 18 "in-situ") carbonate ion concentration. Calcium ion concentrations in seawater vary in direct proportion with the seawater salinity. Becaus e the salinity of open-ocean seaHater only varies within a narro w range, calcium ion concentrations in seawater also vary only within a narrow range Therefore, variation in the concentration of carbonate ion is the primary factor that determines the saturation

PAGE 45

state of seawater with regard to [co3 2-Js is defined as: ( 8) K' sp and as shown below, the quantities !J.s and (1-n) are related: (9a) 6s ([C032-]s [C032-]) [C032-Js [C032-] 6s (9b) (1-Q) ')_ [C03']s [C032-Js It is useful to conceive of these two quantities as "absolute" and "relative" indices of the seawater saturation state. (1-Q) is dimensionless and expresses the degree of undersaturation relative to the exac t saturation point, where Q = 1. Thus, (1-n) may be termed the "relative saturation index" (RSI). In contrast, !J.s is in concentration units (typically sw), and therefore expresses the absolute degree of seawater undersaturation defined by the state of the system and actual ionic concentrations. !J.s could also be called the "absolute saturation index" (ASI). Negative values of !J.s indicate supersaturation, and !J.s = 0 indicates t he saturation point. In the results o f the GEOSECS expeditions (Takahashi et al., 2-2-1980), the quantity DELTA was defined as ([co3 J -[co3 ]s). Thus, DELTA is negative for undersaturation and positive for supersaturation. Calcu lation of the seawater saturation state with a given set of 19

PAGE 46

20 parameters involves equilibrium reactions between different chemical species. Factors which influence the equilibrium are solution composition, temperature and pressure. In order to determine the seawater saturation state, the thermodynamic influence of these factors must be determined precisely. The concentration of dissolved calcium ion in seawater can be calculated directly from salinity. 2+ [Ca ]T, defined as the sum of free and complexed calcium ion concentrations in seawater, is directly proportional (within a variation of 1 .5%) to the salinity, due to the constancy of seawater composition (Culkin and Cox, 1966; Takahashi et al.,. 1 980): ( 1 0) 2.9383 x 104 x Salinity (ppt) This relationship is independent of temperature and pressure. Concentration calculations for the carbonate system components are more complex, because of the interrelationships between the species. The four species in the carbonate system are related in the following manner (Morse and Berner, 1979): (11) + 2-+ + H co3 + 2H The apparent equilibrium constants K 1 and K 2 express the relationship between the hydrogen ion activity, aH, and the dissolved constituents of the system, Hco3-, and co32-

PAGE 47

( 12a) K 1 (12b) K 2 Note that K 1 and K 2 may be expressed on three different scales for hydrogen ion concentration. Morse and Berner (1979) noted that the 21 concentration of hydrogen ion in water "can be defined in several ways depending on which buffer system is used, whether liquid junction potentials are considered, and what definition of ionic strength is used The values of the apparent dissociation constants and calcium carbonate solubility constants are dependent on exactly what definition of seawater pH is used and what standardization technique is used." When expressed in terms of hydrogen ion apparent activity, pH -log aH' refers to the NBS scale, operationally bas ed on defined values of aH for dilute buffer solutions (Bates, 1973; Johnson et al., 1977). + Another scale is the free hydrogen i on concentration, [H ], which i s based on Tris (seawater) buffers (Bates et al., 1977). + Measurements of pH = -log [H ] on the free hydrogen concentration are + expressed as moles H per one kilogram of water. A third scale, that of Hansson (1973), gives pHT = -log[H+]T in terms of the total moles of H+ per kilogram of seawater T hus, three different systems are available for the evaluation of seawater hydrogen ion concentrations. The equations used to calculate the values for K'sp' K 1 ', and K 2 as functions of temperature, pressure, and salinity are delineated in a following section. The relationships described above govern the solution chemistry

PAGE 48

of the seawater carbonate system in a:1 "isolated" sense: that is, not considering the seawater sample as part of the oceanic system. The carbonate ion concentration in a parcel of oceanic water, besides being governed by the influences of temperature, pressure, and salinity, is also influenced by processes which have occurred within the parcel. These processes include the formation or oxidation of organic matter, the injection of atmospheric carbon dioxide, and the dissolution of Caco3 In turn, these processes are related to the history of the water parcel as it has circulated through the oceanic system. The thermodynamic relationships combined with the reactive 22 processes determine the saturation state of the parcel with respect to Caco3 In describing seawater carbonate chemistry it is necessary to first define measurable quantities in seawater that will define the state of the carbonate system of the parcel. A desirable feature of these quantities is that they are invariant with respect to changes in temperature and pressure. Two of the most useful measurable quantities are the total co 2 ( EC02), which quantifies the total inorganic carbon content of seawater, and the total alkalinity (TA). EC02 is defined as: ( 1 3) TA is defined as: ( 14) TA [Hco3-J + 2 [co3 2-J + [B(OH) 4 -] + [OH-] -[H+] + I (weak acid anions alkalinity)

PAGE 49

23 wher-e br-ackets expr-ess the total concentr-ation of each indicated chemical species. The car-bonate alkalinity (CA) is defined as the sum of the fir-st two terms in the TA expression. The carbonate alkalinity and LC02 are related by the equilibrium constants, K 1 and K 2', and the hydr-ogen ion activity, aH, in the following manner: 2 K 'K + K 'a 1 2 1 H ( 15) CA The carbonate alkalinity is not directly measur'ed, but may be calculated by subtracting the borate alkalinity, [B(OH)4 ], from the total alkalinity. The borate alkalinity is calculated fr-om the hydrogen ion activity, the total boron content (BT) of the seawater (which varies directly with salinity), and the boric acid equilibrium constant, K 8': ( 16) BA K + B Variation of the Carbonate System in Oceanic Waters Oceanic processes, whether biological, chemical, or physical in nature, affect the carbonate system in the water column. Changes in the carbonate system will cause the saturation state of the water with

PAGE 50

24 respect to eaco3 to vary. These variations, in turn, affect the distribution of Caco3 sediments on the ocean floor.. Therefore, before discussing the sediment distribution, it is necessary to explain the processes that lead to under-saturation in the water column, and which cause some regions of the ocean to be more undersaturated than others. Caco3 either calcite or aragonite, is more soluble in seawater at low temperatures than at higher temperatures. Either mineral is also more soluble under elevated pressures than at atmospheric pressure. Broecker (1974) shows that in the pressure and temperature range of the oceans, the effect of pressure is more significant than temperature. As temperature decreases from 24C to 2C, at one atmosphere pressure, the carbonate ion saturation concentration for aragonite increases by 18% As pressure increases from 1 to 500 atmospheres, at 2C, the carbonate ion saturation concentration increases by 37%. Therefore transport from warm surface waters to cold abyssal waters will expose a Caco3 shell to more corrosive conditions, simply due to the increased solubility of the mineral under low temperature, high pressure conditions. Broecker (1974) provides a concise explanation of the biological and chemical factors which cause variation in the saturation state of oceanic waters. Because of the simplicity of his explanation, it will be described below. Broecker's process assumes that all water masses will have sufficient oxyge n to facilitate the oxidation of organic matter. Organic matter is continously decomposed by oxidation in a well-circulated water mass. In the deep ocean, organic matter destruction is accompanied by the dissolution of Caco3 These two

PAGE 51

25 processes result in an increase in the and TA content of the water, with increasing more rapidly than TA. In most of the ocean, TA is greater than Oxidation of organic matter therefore decreases the difference between the two quantities. Broecker assumes that this difference is approximately equal to the dissolved 2-carbonate, co 3 content of the water. It should be apparent that organic matter oxidation will reduce the concentration of co 3 2 -in the water mass. decreasing Since ion concentrations are nearly constant, 2-co3 content causes the water mass to be more undersaturated with respect to Caco3 the The circulation pattern of the ocean proceeds by the following generalized route: Water initially sinks fr-om the surface of the northern Atlantic to the ocean floor-, and subsequently flows southwar-d. In the Antar-ctic Ocean, cold water from the surface Antarctic Ocean joins this flow. The circulation then proceeds thr-ough the Indian Ocean to the Pacific Ocean, where the general direction of flow is northward. Tner-efore, the water mass in the deep nor-th Pacific is "aged" relative to the water mass in the northern Atlantic. During the cir-culation, organic matteroxidation and caco 3 dissolution have occurred in the water mass. Thus, deep Pacific water is mor-e undersaturated than deep Atlantic water, even though the conditions of low temper-ature and high pressure are similarin both regions. Pacific water-s are also undersaturated at much shallower depths than Atlantic waters.

PAGE 52

Investigations of CaCOl Dissolution in the Ocean Insight into the nature o f the dissolution process has been gained from the results of several different investigations (Morse, 1983). The primary goal of this research has been to elucidate the factors which affect caco3 dissolution in Investigation of calcite and aragonite dissolution rates in the oceans has involved two methods of research. In-situ investigations have placed caco3 forms in the ocean at various depths. Changes in the forms (i.e. mass loss or indications of corrosion) indicated the 26 magnitude of dissolution. In contrast, laboratory investigations have generally monitored solution effects which accompanied the dissolution of eaco3 Peterson (1966) and Berger (1967) conducted the first experiments that attempted to quantify the dissolution of calcite in the deep ocean. Peterson (1966) suspended machined calcite spheres at various depths in the Central Pacific Ocean. Berger (1967) suspended sample chambers containing assemblages of foraminifera at the same site. At the end of four continuous months of seawater exposure, the spheres and chambers were recovered. Determination of the mass lost by the spheres over the exposure time was accomplished by subsequent weighing. The resulting mass loss vs. depth profile for Peterson's spheres is shown in Figure 3. The figure clearly shows that the dissolution of the spheres increas ed rapidly below 3500 meters. Peterson also noted that the depth at which the rapid increase in dissolution was observed was approximately 1000 meters above the depth of the deepest carbonate sediments in the region.

PAGE 53

27 0 1 E 2 J :t: .._ a. uJ 3 ( Q ' 4 ... 5 0 1 0 2 0 3 0 4 0.5 WEIGHT LOSS (mg/cm2 yr) Figur-e 3 (fr-om Ber-ger-, 1975) Pmfile of dissolution r-ate of optical calcite on a moor-ed buoy in the centr-al Pacific demonstr-ating the existence of a level of r-apid incr-ease of dissolution near-3,700 meter-s Results of Peter-son (1966)

PAGE 54

28 Berger (1967) quantified the effects of dissolution by examining the foraminifera which Pemained in the sample chamber's. He noted appr'eciable dissolution below 1000 meter's, and Papid incr'eases in the degt:'ee of dissolution below 3000 and 5000 meter's. Dissolution caused changes in the species composition of the assemblages, the size distribution, the number' of appar'ently "damaged" or' par'tial shells, and the aver'age particle weight. Berger states that his experiments, along with the investigation of Peterson, ver'ified hydrographic and thermodynamic predictions of undersaturation with respect to calcium car'bonate in the water column. The marked changes observed at and below 5000 meters were definitely due to the "pPonounced undersatur'ation of the abyssal water." Berger's statement connects the thermodynamic state of seawater (undepsaturation) with the kinetics of calcite dissolution. Note that a var'iation in thermodynamic state does not necessar'ily imply a connection with the kinetics of a chemical pr'ocess. A subsequent investigation of the pr'eset:'vation of foraminifera in sediments (BePger', 1968) cot:'r'elated the Pesults obser'ved in the moo Ping experiments. BePger' devised a "solution index" based on species-selective dissolution in Ot:'der' to quantify the effects of dissolution on capbonate sediments. As shown in Figur'e 4, incPeases in the number of "resistant" for'aminifer'a relative to the number of "easily dissolved" fomminifer'a indicated when dissolution began to substantially affect the composition of the sediment. Note that the lar'gest change in the "solution index" in Figur'e 4, for' the Atlantic Ocean, is approximately 1 000 meter's deeper' than the depth of the transition to increased mass loss in Figur'e 3, for' the Pacific.

PAGE 55

29 0 1 2 I 3 I -E .X :c t-I l&J 0 \ 4 .\ 7 5 ., I 0 6 0.8 1 0 1.2 1.4 SOLUTION INDEX Figure 4 (from Berger 1975) Profile of the preservati o n aspect ("solution ind ex ) of foraminifera in the central Atlantic, demonstrating an abrupt transition from well-preserved to poorly-preserved foram assemblages ( foram "lysocline"). Points plotted are averages (mean N = 7). Results of Berger ( 1 968).

PAGE 56

30 In a later article by Berger (1975), the author noted that the initial investigation of foraminiferal preservation had led to a entire body of research concerning "preservation facies", and their chemical and paleo-historic implications. In another moored experiment, Milliman (1977) suspended aragonitic ooids in the Sargasso Sea and investigated the amount of mass loss in a manner similar to that of Peterson ( 1966). He observed mass loss and erosion of the ooids, and noted a pattern of increased dissolution with depth similar to that of Peterson. However, .his results indicated slower dissolution than other (prior and subsequent) investigations. Possible explanations for this discrepancy are the nature of the materials used, and partial saturation of the water which became entrained in the sample chamber. Honjo and Erez (1978) also performed an in-situ experiment that utilized chambers containing biogenic calcite and aragonite. Both synthetic aragonite and pteropod shells were used. In order to reduce the effects of saturation within the chamber seawater was pumped through the chambers at a constant rate. The rates observed in this experiment were also slower than rates determined in laboratory investigations, but showed the expected pattern of increasing rate with depth. However, the ratio of the pteropod dissolution rates to the synthetic aragonite dissolution rates showed substantial variability, which may have been caused in part by the large differences in surface area between pteropod shells and synthetic aragonite. The authors also indicate that pumping did not fully mix the caco3 materials and seawater. Incomplete mixing caused the chamber contents to collect at the bottom of the chamber ( "pending"),

PAGE 57

31 which could inhibit the dissolution of the material (Morse et al., 1979). Another investigation of in-situ dissolution was performed in the Panama Basin, and utilized foraminiferal tests as the experimental material (Thunell et al., 1981). In this experiment, weight loss and broken tests were used as indicators of dissolution. The researchers found that accelerated dissolution began below 2750 meters. The researchers noted that the sedimentary and hydrographic lysoclines were not found at the same depth as the saturation -undersaturation transition depth in the region. Laboratory Investigations of caco1 Dissolution The alternative to in-situ experiments is laboratory experimentation. One of the most important advances in the laboratory investigation of carbonate dissolution, as well as the dissolution or precipitation of many other minerals, was the application of a "pH-stat". The "pH-stat" is a version of a "chemostat". The operation of the chemostat enables the maintenance of a constant solution composition while reaction takes place. The pH-stat monitors the dissolution rate of a mineral by measuring the rate of reagent (generally acid or base) addition required to maintain a constant solution pH. The use of the "pH-stat" in the investigation of calcite dissolution kinetics was first described by Morse ( 1974a). Several subsequent investigations of Caco3 dissolution (Morse et al., 1979; Keir, 1980; Walter and Morse, 1985) have employed versions of the pH-stat. According to Morse

PAGE 58

32 ( 1983), "the constant composition approach has proven to be particularly useful in studying a number of problems where extensive reaction is required (e .g. inhibitor influence on reaction rates, and coprecipitation reactions) and where factors other than saturation state are being studied. However, an alternative method for the investigation of mineral dissolution also exists. This method is called the "free-drift" method, in whic h the composition of a solution sealed within a reaction chamber is allowed to vary as the reaction proceeds (Weyl, 1965). The "fr-ee-drift" method has been used widely to investigate both r-eaction kinetics and the solubilit y behavior of many miner-als. Two of the main deficiencies of the "free-drift" method are the amount of time necessar-y to induce measurable changes in the composition of the contained solution, and the variation in reaction rate or chemical behavior that accompanies alter-ation of the solution composition Also, at high r-ates of reaction, the "pH-stat" vtas not able to maintain a condition of "dynamic equilibr-ium" (steady-state reaction), where the r-ate of acid addition balances the rate of caco3 dissolution. The inability to maintain a constant steady-state could have resulted in the deter-mination of a false dissolution rate (Mor-se, 1983). The first discussion of Caco3 dissolution kinetics in which provided a theor-et ical basis for the dissolution process on the molecularlevel, was presented by Berner and Mor-se ( 1974) Their investigations wer-e preceded by several deter-minations of calcite and ar-agonite solubilities in seawater (Weyl, 1965; Macintyre, 1 965; Ben-Yaakov et al. 1971, 1974; Ingle et al., 1973). Berner and Mor-se

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found three separate regions of dissolution behavior (Figure 5) with increasingly undersaturated conditions. However, the most important region relative to oceanic conditions was the near-equilibrium region "3". In this region, dissolution rates increase very slowly with increasing undersaturation until a critical level of undersaturation is reached. At this critical point, the calcite dissolution rate begins to increase more rapidly with increasing undersaturation. The critical undersaturation point was influenced by the concentration of dissolved phosphate. Higher concentrations of phosphate moved the critical level of undersaturation further from the equilibrium saturation point. Morse and Berner stated that the transition from slow to rapid dissolution corresponded to the degree of undersaturation in the oceans which produced the lysocline. The experimental material used was a fine-grained synthetic calcite. 33 Walter and Burton (1986) confirmed that phosphate inhibited the dissolution rates of calcite and aragonite. Inhibition increased by a factor of two as Q (i.e. the degree of undersaturation) increased from 0.1 to 0.8. However, Walter and Burton conducted their experiments between 7.0 and 7.5 pH. Sjoberg (1978) reported a lesser degree of rate inhibition due to phosphate at pH 8. My experiments were conducted in seawater with phosphate concentrations ranging from 0.3 to 4 )..llTIOl/liter. Phosphate was determined by Auto-Analyzer on research cruises. Phosphate concentrations in the seawater used in the laboratory were analyzed colorimetrically using the method of Strickland and Parsons (1972). These concentrations are well below the concentration of 50 llmol/liter used in the Walter and Burton study. The authors stated that concentrations of phosphate lower than

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34 t a a::: ApH-Figure 5 (from Morse and Berner, 1974) Schematic plot of rate of dissolution versus The inset is an enlargement of the region near equilibrium demonstrating the presence of a pronounced discontinuity which is otherwise lost on the scale of the main diagram. This discontinuity is believed to represent the chemica l lysocline.

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50 IJlllOl/liter provided results indistinguishable from studies in low-phosphate (less than 1 seawater. Morse (1978) described the dissolution of calcium carbonate in deep-sea sediments, and expanded the results of Morse and Berner (1974). Morse used an equation which e x pressed the dissolution rates as a function of the quantity n (Equation 7) of this form: ( 17) Rate (% of caco3 mass dissolved/day) 35 An investigation of aragonite dissolution kinetics using both synthetic and biogenic aragonite was performed by Morse, deKanel, and Harris (1979). The pteropod shells were taken from a deep-sea core. The investigators found that two rate laws described their data, with a transition occurring at a value of n = 0.44. The two functions which describe the data are: ( 18) :: Q 0.44 Rate (% I day) ( 19) 0.44 Q ;;: 0.25 Rate ( % I day) 1 31 8 ( 1 -Q) 7 27 Only one point was generated with the pteropod shells. This point indicated a slower dissolution r ate than the corresponding dissolution rate of synthetic aragonite (Figure 6). The reduced surface area of the pteropod shells relative to the synthetic aragonite was cited as the cause of the reduced rate. An investigation of biogenic calcite and aragonite dissolution kinetics in seawater that utilized a variety of materials was

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36 ...._Pteropods 0 Synthetic ArOCJOnit e ... .c (,) 0 e -2 0 Cl: 0 0 -4 -1.2 0 8 -0. 4 0 loo(I-D.) Figure 6 (from Morse et. al., 1979) The log of the dissolution rate versus the log of (1-Q) for synthetic aragonite and pteropods Note the single point for pteropod shells, and the substantially reduced dissolution rate of the shells.

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performed by Keir (1980) Keir used synthetic calcite and aragonite, various species of foraminifera and coccoliths, individual pteropods from a marine sediment sample, and a sediment sample from the Ontong-Java plateau. Keir found that the value of the exponent in equation 17 was 4.5 for calcite and 4 2 for aragonite, and the best-fit equation for the dissolution of pteropod shells was: (20) Rate (% I day) 318 (1 n)4 2 37 Walter and Morse (1985) used the pH-stat method to investigate the dissolution kinetics of shallow marine carbonates, composed of both calcite and aragonite. The reaction order (value of n) for these e x periments was found to be generally lower than the reaction orders reported prev iously. The reaction order low-Mg calcites was near 3, and t h e for the various forms were approximately 2.5. The dissolution equation the gastropod Strombus is given below: ( 21 ) Rate (% I day) 302 (1-n)2 54 A comparison of the results from et al., Keir, and and Morse is shown in Table 1. Other investigations of calcite dissolution have been by Plummer et al. (1978; 1979), and Chou (1987). The methods and solution compositions in these experiments differed markedly from my investigations and the investigations discussed the of Plummer and Chou will not be here, but the

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TABLE 1 Comparison of aragonite dissolution rate predictions from equations of Morse et. al. (1979), Keir (1980), and Walter and Morse (1985) Dissolution rates are expressed as percent of aragonite mass dissolved per day. [co3 2-Js = 6 .469 x 10-5 mmol/kg sw1 or [co3 2-Js = 7.597 x 105 mmol/kg sw2 Morse et. al. Keir Walter and Morse Eqn. 202 Eqn. 211 6 .144 0.86 0.30 0 .05 5.820 1.6 0. 71 0 .35 5.497 2 6 1.4 1 2 mm:::>l 5 .174 3 9 2.6 5.1 per 4.850 5 6 4.5 9.0 kg sw 4.527 7.7 7.1 14.2 X 4.204 10.3 10.7 21. 0 105 3.880 13.5 15.8 29. 5 3.557 17. 3 22. 4 39.7 3.233 21 .6 (23.3) 3 30.9 51.9 1Calculated with K' = 6.65 x 107 mol2 (kg sw)2 For seawater spa 2+ of 35/00 salinity, 25C, and atmospheric pressure. [Ca J = 1.028 x 102 mol/kg sw. 2Calculated with K' = 7.81 x 107 mol2 (kg sw)2 spa 3Based on Eqn. 19. 38

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39 relationship of their research to our results will be presented in the final chapter of the dissertation. It is clear that the dissolution of Caco3 is an important oceanic process, and has been an active area for research. However, several topics were not addressed in previous investigations. The topics that formed the basis of the investigation described in this dissertation are presented below. Investigations of carbonate dissolution kinetics in the laboratory are removed from the oceanic environment, and therefore are only simulations of the dissolution process in the oceans. The fonn of naturally-produced minerals may be somewhat different than the materials which are used in laboratory investigations. If the dissolution characteristics of "fresh" pteropod shells differ from synthetic aragonite or shells taken from cores, it would be desirable to determine the nature of that difference. Furthermore, many laboratory investigations have taken place at room temperature and atmospheric pressure. Some investigations have inc luded the influence of temperature (Plummer, 1978; Mucci, 1983), but rarely the influence of pressure. Since most of the dissolution 0 of Caco3 in the ocean takes place at temperatures close to 2 C and at very high pressures (greater than 500 atmospheres), observation of dissolution under these conditions is also desirable. Based on a qualitative comparison, a basic philosophical difference between laboratory and in-situ methods is apparent. The in-situ methods provide insight into the actual workings of the oceanic system despite difficulties in precise measurement and experimental control. The laboratory methods enable precise,

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controlled experimentation in an setting that is one or more steps removed from the oceanic environment. 40 Therefore, a method of investigation that incorporates the best features of the "in-situ" and "laboratory" approaches is a desirable goal. In order to advance the investigation of carbonate dissolution, a method was developed for this research with several new features which allowed new insight into the aragonite dissolution process. These features included: a) high sensitivity to compositional changes in the solution; b) rapid conduct of experiments; c) utilization of natural samples; d) transportability; and most importantly, the system was e) capable of simulating oceanic conditions of temperature and pressure in a laboratory setting. The overall goal of the research is to utilize this system to gain further insight into the relationship between ocean chemistry and the calcium carbonate minerals produ ced by oceanic organisms. The primary goal of this dissertation is to delineate the factors and processes involved in the dissolution kinetics of aragonite in the deep ocean. In conjunction with this investigation, an estimate of aragonite and calcite fluxes and the magnitude of aragonite dissolution i n the oceanic system will be included.

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CHAPTER 2: Caco3 FLUX RESEARCH METHODS and RESULTS "Hast thou entered into the springs of the sea? or hast thou walked in the search of the Depth?" The Bible, Job 38:16, KJV. The inspiration for this body of research may be found in The Fate of Fossil Fuel co 2 in the Oceans, edited by Neil R. Andersen and Alexander Malahoff. In the paper entitled "Sedimentation and Dissolution of Pteropods in the Ocean" (Berner, 1977), the author states: "Three methods can be used to estimate the fraction of calcium carbonate delivered to the sea floor as aragonite, rather than as calcite. The first, and most direct, is the quantitative mineralogical analysis of falling particles collected on sediment traps suspended at various depths in the ocean To the writer's knowledge such a study has not been undertaken and is sorely needed." Although an initial study similar to the one described above was subsequently performed (Berner and Honjo, 1981), the necessity for an extensive study of this topic still remained. Researchers from the University of South Florida and the Pacific Marine Environmental Laboratory collaborated on three cruises in the North Pacific Ocean: one during the summer of 1981 (on board the National Oceanic and Atmospheric Administration vessel R / V Miller Freeman, prior to the 41

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author's participation); and another during the summer of 1982 (on board the NOAA vessel R/V Discoverer, Figure 7). These cruises included the deployment of free-floating sediment traps at various depths for relatively short periods of time. Detailed hydrographic work concerning the oceanic saturation state along each cruise track was performed, along with pressurized dissolution experiments using pteropod shells from the sediment traps. On a third cruise, multiple sediment trap deployments were made at two stations in the North Pacific Ocean in the summer of 1985 from the R/V Discoverer (Figure 7) Following on-shore analysis of the sediment trap collections and the data from the dissolution experiments, a new experimental method for the investigation of aragonite dissolution under pressure was developed. This method was subseq uently employed on two research cruises, in the southern Indian Ocean dur ing the spring of 1985 (on board the Territoires Australes et Antarctiques Francais (T.A.A.F ) vessel R/V Marien-Dufresne, Figure 8), and in the North Pacific Ocean during the summer of 1985 (on board the R/V Discoverer). After the cruises and a data-analysis interlude, a slightly modified version of the experimenta l method used at sea was employed in the laboratory. 42 The research effort has been noticeably concentrated in the North Pacific Ocean. Factors which make this oceanic region of special oceanographic interest are the shallow depths of the saturation horizons for calcite and aragonite, and the high degree of undersaturation of the water mass beneath the saturation horizons. The sediment trap effort was vital as both inspiration for the initial dissolution study, and the traps also served as a source of

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8-11 12 8 6 5 ... o HAWAIIAN IS. 3 KWAJALEJN Figure 7 -Cruise tracks and stations for Nort h Pacific cruises on board the R/V Dis coverer. a. May-June 1982 b June-July 1985. 43

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c( (.) a: u. c( ::1 < 0 a: w .... <1.1 ::1 < D .... w N 0 a: u Figure 8 -Cruise track of the R / V Marien-Dufresne in the southern Indian Ocean and circumpolar Antarctic Ocean (INDIGO I cruise), February-March 1985. 44

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naturally-occurring pteropod shells. The methodology of the sediment trap research is described below, along with a short discussion of the initial results. A description of our dissolution rate determinations will be presented in the following chapter. Sediment Trap Experiments Summer 1982; Summer 1985 The vertical fluxes of particulate matter in the North Pacific Ocean were estimated with free-drifting sediment traps (cross sectional area, 0 .66 m 2 ) patterned after the cone design of Soutar et al. (1977). (Figure 9) In this system, trapped materials are concentrated in a Teflon receptacle (2.5 em in diameter) at the bottom of the trap's cone. Prior to retrieval, the Teflon cup and its contents were isolated from further inputs by a double ball-valve seal activated by an electronic release. The total particle flux was determined gravimetrically. Immediately after recovery, particulate materials from the respective traps were filtered onto tared (47 mrn in diameter) Nuclepore membranes and then briefly rinsed with distilled-deionized water. The filters were stored in a vacuum desiccator with silica gel until they could be weighed on a microbalance in a shore-based clean facility. Pteropods and foraminifera were identified microscopically, removed from the samples, and transferred to tared Nuclepore membranes with a small nylon brush. To minimize sample contamination the microscopic examination and transfers were made in a vertical-flow clean bench. After desiccation under vacuum, the biogenic materials were weighed on a microbalance housed in a clean room.

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FIBERGLASS COWLING Js-1 1 WAY BALL CHECK VALVE PVC SUPPORT I ELECTRONI C TIMER-I l j I' ./ , I' I' I' I 1-STAINLESS STEEL BRACKET -PVC SUPPORT LEG NICHROME BURN WIRE DOUBLE BALL VALVE SEAL SYSTEM -PVC BASE PLATE 46 Figure 9 (from Betzer et. al. 1984a) Sediment trap of the design used by Betzer et. al. (1984b) for North Pacific cruises on-board the R/V Discoverer

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Sediment trap deployments were made at seven stations in the western North Pacific in the summer of 1982, for collection periods not in excess of 40 hours. This short deployment period reduced the possible loss of aragonitic material due to dissolution, which may have occurred in sediment traps deployed for longer periods of time. 47 Several large pteropods (species incl. Diacria trispinosa, Cuvierina columnella, Cavolinia tridentata, and Clio pyramidata) were excluded from the flux calculations, as these species are both large and rapid swimmers, and were believed capable of actively entering the sediment traps. The flux rates were calculated on the basis of the trap's collection area (0.66 m 2), and the length of time the trap was deployed. (Deployment times ranged from 28 to 40 hours, covering at least one diel cycle). The aragonite flux was calculated primarily from the masses of small pteropod species, principally Limacina bulimoides, inflata, and helicina. Large, empty shells, such as a specimen of Clio pyramidata found in a trap deployed at 2200 meters, were included in the flux calculation. Figure 10 (Betzer et al., 1984) shows the fluxes calculated for the seven sediment trap stations. The figure also shows the 100%, 80%, and 60% saturation horizons (based on n calculations) in the region. The figure shows that the aragonite fluxes are substantially reduced from 100 meters to 2200 meters. There is a substantial reduction in the flux between 100 and 400 meters, above the aragonite saturation horizon. This reduction may be due to transport of pteropods by vertically migrating predators, as suggested in Betzer et al., 1984), or biological interactions with the traps themselves, which will be discussed in a later section. These interactions are

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Sta t ion 3 5 6 8 1 2 (Kuroshio Station) 18 20 5 0 32 6 16 0 29 100 400 ...... E ..... J: Q. 900 1 4 1 7 1 3 0 6 0 -\ 1 \ \ \ 2200 { \ I } 1 9 1 .1 0 7 0.6 0 2 I I I I I I I I 12 18 24 30 36 42 48 54 Latitude (0N) Figure 10 (from Betzer et. al. 1984b) Depth distributions of pteropod fluxes (in milligrams per square meter per day) from the western North Pacific. Aragonite saturation state is shown for 100 percent saturation ( ) 80 percent saturation ( ----), and 60 per cent saturation(-). Calculated f l uxes do not include 10 perce n t ) contributions from pteropod fragments 4 8

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due to the fact that pteropod abundances are high at 100 meters, and some organisms may have actively entered the traps. 49 The data in Figure 10, combined with measurements of the foraminiferal flux, indicated that aragonite constituted a larger percentage of the oceanic Caco3 flux than had been shown in previous sediment trap studies (Berner and Honjo, 1981). The primary reason for the difference was believed to be the dissolution of aragonite in traps which had been deployed for extended periods of time in undersaturated seawater. Under such conditions, aragonite would be lost to dissolution, and the measured flux of calcite would therefore be enriched relative to aragonite. However, the short (less than 1.5 day) deployment times used by Betzer et al. limited the amount of aragonite which would be lost to dissolution. Long-term sediment trap deployments have ranged from 75 to 200 days. In these studies, the majority of aragonite collected in traps which were deployed in deep, undersaturated waters may have been subsequently dissolved. It is also apparent in Figure 1 0 that the flux between 100 meters and 400 meters was also substantially reduced, in the shallow region of the water column which was supersaturated with respect to aragonite. As dissolution would not influence the trap's contents in this region, biological behavior such as evasion or predation may have contributed to the alteration of the aragonite flux.

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CHAPTER 3: METHODS OF DISSOLUTION RESEARCH "He gathers the waters of the sea into jars/ He puts the deep into storehouses." The Bib l e Psalms 33: 7 NIV. 50 Due to the variety and evolution of the experimental methods used in our dissolution research, the methodology employed in each "cruise phase" will be discussed separately. The procedural modifications made in the laboratory and the calculation algorithms will also be presented in this chapter. Dissolution Methods -Summer 1982 The dissolution experiments performed o n the R / V Discoverer during the summer of 1982 employed the f ollow ing methodology Individual shells of pteropods that were devoid of internal organic matter were used. No method was used to remove organic coatings, as the reaction rate of the natural surface was desired. In some cases, the shells wer e collected empty. If, however the body of the organism was still present in the shell, the body was carefully removed prior to the use of the shells in experiments Microscopic examination and flushing of the shell with seawater (using a syringe needle) insured complete removal of all visible organic matter. The shells were housed in glass-plastic sample vials, which were placed in thermostated high-pressure vessels. The vesse l s were gently oscillated at seven cycles per minute for the duration of each

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51 experiment. All experiments were conducted at 5C and at pressures near 375 atm; the extent of dissolution was monitored by measuring the seawater pH in the sample vials before and after each experiment. The pH changes i nduced by caco 3(s) dissolution (22) + 2 + + H Ca were generally quite small (0 .03 pH 0 20) In order to maximize the signal-to-noise ratio, our seawater sample volumes were small(-9 ml) and pH was measured using electrodes (Ross type reference Orion Research) substantially free of drift in the course of thermal cycling. The total Caco3 dissolved in each experiment was determined using the initial alkalinity, pH, and total co 2 characteristics of our seawater samples, and by calculating the alkalinity change (twice the change in total dissolved co 2 ) necessary to produce the observed change in pH. The equations and f o rmation constants used in our calcul3tions follow Millero (1979), Feely et al. (1984), and Byrne et al. (1984). Feely et (1984) describes the measurement of TA and Subsequent measurements of sample ( pteropod) masses on laboratory microbalances permitted the result of each experiment to be expressed in terms of a % weight loss. Feely et al. (1984) uses mercuric chloride (HgC12 ) to limit bacterial respiration in his seawater samples prior to analysis. Earlier experimentation by Byrne and associates on-board the R/V Miller Freeman indicated evidence of bacterial respiration when fresh pteropod shells. Therefore, for the potentiometric experiments only,

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52 mercuric chloride was added to the solutions, so that Hg( II) in solution would remedy the possibility of bacterial respiration in the dissolution experiments. Our seawater media were 2 x 10-5 M in Hg(II). Hg(II) has a very substantial affinity for halide ions, so Hg2+ i s not readily available for interaction with carbonate surfaces. 2-HgC14 substantially dominates the speciation scheme of Hg(II) in t d d f Hg2 + -1 9 seawa er an 1n our me 1a, ree concentrations are -10 M. Analysis Algorithm for Potentiometric Data The cumulative % dissolution in each of the potentiometric experiments was modelled by simulating the dissolution process in the vial with a computer program. The program integrated the dissolution process over the length of the experiment, taking into account the deceleration of the dissolution rate as the solution became increasingly saturated. The program generated a predicted value for the total amount of dissolution, D, using the following integral form of the classic rate equation, equation 17 (Morse, 1978): (23) D Dissolution times in these experiments ranged between one hour and one day The integrations were performed numerically by dividing the total dissolution time, t, in each experiment, into minute-long segments. Our analyses indicated that dissolution segments 1 minute produced identical model predictions. Equation 23 -was applied to each of our 80 dissolution experiments In each analysis, a set of parameters (k, K'spa and n)

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53 was chosen and held constant. Our choices of K were con strained spa by the results obtained i n a variety of aragonite solubility determinations (see Appendix I). As described above this process generated a value, D(predicted), for each of our experiments. D(observed) was calculated from the pH change of each experiment. The differences,
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Figure 11 Illustration of high-pressure chamber and variable pathlength spectrophotometric cell (shown with ) used in this study.

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55

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pressurization (Byrne, 1984). In my experiments, pteropod shells and seawater containing the pH-sensitive dye phenol red at a concentration of 8 x 10-6 M were placed within the variable pathlength cell. (The seawater was not treated with HgCl2 for the spectrophotometric experiments.) The cell was placed within a thermostatted outer pressure housing, which was closed with 3/4-inch pressure-bearing plexiglas windows (Figure 11). 56 Following pressurization, absorbance readings were taken with a Varian Instruments DMS-90 spectrophotometer. In initial "feasibility" experiments, the thermostatted pressure chamber which contained the variable pathlength spectrophotometric cell was oscillated by hand. A primary goal of this initial stage of research was to demonstrate that the system worked as expected. Over a period of time in a "free-drift" experiment, the reaction rate should diminish due to the increasing saturation state of the contained seawater sample, caused by the dissolution of an aragonite shell. Figure 12 shows the results of a four-hour experiment which demonstrated that saturation does occur, and also demonstrated that experiments with a duration of one hour or less essentially measured initial dissolution rates. Note that absorbance measurements in Figure 12 were spaced approximately four minutes apart. This spacing was found to cause a slight inhibitory effect on the observed rate of reaction. The inhibitory effect was ascribed to "ponding", which is the formation of a saturated boundary layer at the surface of the shell, which protects the shell from the corrosive effects of the bulk solution. When the measurements are spaced closely together, the shell is immobile for a

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. 12 effects of solution on dissolution of Note that the 60-85 minutes of is equivalent to initial dissolution

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w (_) z <{ co 0: 0 (/) co <{ ABSORBANCE vs. TIME ARAGONITE DISSOLUTION . . .. 0 30 60 90 120 150 180 210 240 270 TIME (MINUTES) .q(\J 0 0 II I 0.
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59 large of the time, which allows the formation of the the between the length of time the shell was immobile, and also the change in the signal. An of 20 minutes between was used in typical which below. (Dissolution inhibition due to incomplete mixing is a of The effect influences dissolution but is not a which will influence dissolution conditions.) My initial indicated that exchange of within the and within the cell could be eliminated by the "fit" of the on the cell windows. fit windows exhibited no exchange of solution when a strongly-colored acid fuchsin dye solution was used to fill the chamber, even when ized in excess of 5000 p. s. i. Fortunately, this allowed the use of distilled to fill the pressure chamber, which limited chamber corrosion. The housing containing the cell was oscillated at approximately 6 cycles/minute between readings (manually on the mechanically on the and in the laboratory). The cell and its contents maintained at 5 + 0.2C with a Lauda K2-R Absorbance measurements were made after an initial 1 0 -minute equilibration period following the and at subsequent 20-minute intervals. In most 20-minute intervals sufficient to provide a change. The dissolution rate was calculated over the hour which encompassed the

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three 20-minute intervals, and did not include absorbance changes which occurred in the initial 10 minutes. 60 Each absorbance measurement included absorbance readings obtained at 558 and 700 nm. As phenol red absorbance at 700 nm is negligible, readings at this wavelength allowed normalization for slight variations in cell alignment. Measurements of a "blank" seawater sample (seawater without dye) enabled instrumental baseline corrections. All of the absorbance readings taken on board the Discoverer were signal-averaged with an Apple IIe computer interfaced with the spectrophotometer, which increased the sensitivity of the technique and compensated for signal variations caused by the motion of the ship. (The signal-averaging program was not available on-board the Marien-Dufresne, so the data were recorded by hand.) Calculations Defining the o ceanic carbonate system on the basis of measurable parameters and thermodynamic relationships is a continuing process. This research has relied on the results of several investigations, which strove to define the variability of carbonate dissociation constants and solubility products as functions of temper-atur-e, pressure and salinity. A par-ticularly cr-ucial step in this r-esearch has been choosing the most r-eliable set of r-elationships which define the carbonate system. Feely et al. (1984) describes the data collection and analysis techniques that were used on board the R/V Discovererin the summer of 1982. Water samples were collected in 30-li terNiskin bottles and

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61 immediately transferred into 1-liter glass-stoppered bottles containing 1 .0 ml of a saturated solution of HgC12 to decrease bacterial oxidation of organic matter prior to analysis. The samples were stored in a dark cold storage room at 4C for as much as 1 2 hours. The samples were analyzed by a Gran potentiometric titration using a Brinkmann E636 titroprocessor linked to a Hewlett-Packard 35 computer. The data from the t i troprocessor were automatically fed into the computer and processed using the modified Gran equations described by Bradshaw et al. (1981). Corrections for boric, silicic and phosphoric acid were computed from equations similar to Takahashi et .al. (1982) in the GEOSECS Pacific Expedition report. Total borate concentration was computed using the relation given by Culkin (1965). The dissociation constants of carbonic acid and boric acid for the titration analysis are from the work of Almgren et al. (1977). Feely et al. utilized the equations of Millero (1979) for the calculation of the water column saturation state. Millero's equations were determined by first evaluating several sets of data for the dissociation constants of carbonic acid and boric acid in pure water, and then using a least-squares computer program to fit the data sets to adjustable parameters in an equation of the following form: (25) A + B/T + C l n T T is the absolute (Kelvin) temperature. The salinity dependence of the dissociation constants in seawater is expressed by an equation of this form:

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(26) ln K. I 1 wher-e Ki is gener-ated by the equation for-temper-atur-e dependence (equation 25), and Sis the salinity in par-ts perthousand (0 /00). 62 The calculation forthe pr-essur-e dependence of the dissociation constants is slightly more complex. Miller-o (1979) selected the following equation to express pr-essur-e dependence: (27) + (0 .5 and are the molarvolume and molar compr-essibility (the change in the molar volume with pressure) for a r-eact ion. R is the gas p constant, and Ki0 and Ki ar-e the values of the dissociation constant at atmospher-ic pr-essur-e and at a higher pr-essure P, r-espectively. Subsequent resear-ch by Mucci et al. ( 1 982) on the pressur-e dependence of the aragonite solubility pr-oduct indicated that the compressibility term was unnecessary. (However, the compressibility ter-m r-emains in the equations used to calculate the pr-essure dependence of K1 ', K2', and The pr-essur-e dependence of K' is then dir-ectly calculated to the molar volume change forspa dissolution. Feely et al. (1984) used the equation from Mucci et al. (1982), which we have also used in ourcalculations: (28) ln (K' p /K' 0 ) sp sp Our method of calculating the temperatur-e and salinity dependence of the solubility product of aragonite, K'spa' utilizes the equation

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63 from Ingle (1975) for the temperature and salinity dependence of the calcite solubility product, K' However, Ingle's equation was spc modified so that the expression provided the solubility product of aragonite determined by Morse et al. ( 1980) and Mucci (1983). Ingle's equation with the new coefficients used in our work is: (29) K' spa [-49.842-57.68 S1 3 + 159.4 2 logs 1.096 X 105 T2J 10-7 This expression gives a va lue of K' essentially identical to the spa value reported by Morse et al. (1980) and Mucci ( 1983), which is 6.65 x 10-7 moles 2 kg2 at 25C and 35 00 salinity. Given these same conditions, Equation 29 predicts a value of 6.66 x 10-7 moles 2 kg2 for the apparent solubility product. Feely et al. ( 1984) used the equation for the temperature and salinity dependence of the aragonite solubility product given in Mucci (1983). Data provided by Feely for the Eastern North Pacific water colwm can be used to compare the results of the two calculation methods. At 50N, 145W in July 1985, the value of n at 350 meters calculated by Feely was exactly equa l to 0.6. Our method produ ces a value of 0.584 for the same data. The ratio of the two quantities, 0.6/0.584, is 1 .028. This difference results from the slightly higher solubility product for aragonite which is calculated by our equation as compared to Mucci's equation. The equation used in this work produces a value of K' identical to Mucci's value at 25C and 35 oo salinity. spa However, Mucci's equation produces a value for K'spa of 6.48 x 107

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moles 2 kg2 for the same conditions. Note that the ratio of the two solubility products is 6.65/6.48, or 1.028, which is exactly equal to the ratio of the water column saturation state determined above. The reason for the difference between the two values of K' may lie in spa the fact that Mucci's salinity dependence was determined for a salinity range of 5 to 44 /00, and Ingle's equation was for a range of 27 to 43 /00 Therefore, Ingle's equation may be more suitable for the salinity range of 33 to 36 /00 The slightly smaller value 64 of n which is calculated by our method stems from the greater value of K' which is produced by my equation. spa A more detailed discussion of the functions used to calculate the values of K' and K' appears in Appendix 1 spc spa

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65 Analysis Each data set a the initial total dioxide initial total alkalinity ( TA)i' and initial NBS scale pH (paH)i the sample used in the measurements calculation of 6pH time using the following equation et al., 1985): (30) 6pH ( A 2 A ) log mm (Amax (A1 -A ) log mm (Amax-A1) In this equation, A 1 denotes an initial measurement and A 2 denotes a subsequent absorbance A and A are max m1n the and bound values, at high and low pH, the same concentration of phenol red. NBS scale p H (paH = -log aH) is then calculated as: (31) + 6pH Dissolution induced changes in total alkalinity and total co2 are the equation (32 ) 6TA The total alkalinity in our closed system can be written as (33) (TA\ + 6TA = I I I 2 K 1 K 2 + K 1aH I I I 2 K1K2+K1aH+aH I I I BT is the total boron concentration and K 1 K 2 K8 are carbonic and boric acid dissociation constants calculated according to the methods of Millero (1979). By combining equation (32) and (33), 6C02 can be calculated through time as:

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(34) where (TA)i -Rc (EC02)i R 8 BT R 2 c 66 Our 6C02 measurements were subsequently converted to a dissolution rate in percent per day using the mass of the pteropod shells, which were dried and weighed to 0.1 the cruise. The surface area of the materials used in the experiments was to be held through the use of single pteropod shells in each set of experiments. Least-Squares Curve Fitting Technique Severa l datasets of dissolution rate versus 6s or (1-n) were analyzed using a least-squares curve fitting technique. The program makes use of the Marquardt NLIN iteration technique for fitting data to a non-linear model. The Marquardt technique is available from the Statistical Analysis System (SAS) package, which was accessed for this research through the University of South Florida mainframe computer system. (The statistics package was created by the SAS Institute, Inc.) The technique essentially minimizes the residual sum-of-squares, L where = D(observed) D(predicted), and D is the value of 1 1 the dependent variable for each data point. For the data generated in this study, predicted (model) dissolution rates were provided by the expression

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67 (35) R = Ax.n In this equation, R represents dissolution rate, X i s saturation index or 1-n), and A and n are constants whose value is determined by the least-squares analysis of the data. The program also generates asymptotic standard errors and 95% con fidence intervals for the variables A and n. Equation 17 is in the form of Equation 35, where x corresponds to the saturation index (1-n). For clarity, the equation using the saturation is of the following form: (36) Rate(% /day) K ([co32-Js-[co3 2-])n Asthe functional form of Eqns. (17) and (36) is equivalent in form to Eqn. (35) data were analyzed by the least-squares method in an identical manner. In geometrical terms, the program generate s a residual sum-of-squares value, which is calculated as the sum of the squared vertical distances of the data points from the model curve. The lowest residual sum-of-squares indicates the best curve Hhich can be fit to the data points. The major application of the residual sum-of-squares values in this research has been to assess which saturation parameter, or (1-n), is better suited to describe the relationship between aragonite dissolution rates (R) and the seawater saturation state. A schematic diagram of the entire experimental technique, from initial data collection to least-squares statistical analysis, is shown in Figure 13.

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INITIAL DATA COLLECTION PDP l J ffiiDillllPUTEB FOR DATA STORAGE, ANALYSIS AND COMPILATION IB!\ 1'\AINFRA/'\E SY STE/'\ FOR STATISTICAL ANALYSIS OF DATA Figur"e 13 Schematic diagr"am of spectr"ophotometr"ic analysis technique. Initial data was collected by spectr"ophotometer" inter"faced with Apple IIe computer'. Water" samples wer"e collected by bottle cast Or' fr"om seawater" samples in laborator"y, analyzed by titr"ation. Data from exper"iments was transcribed and filed into minicomputer" and analyzed by custom pr"ogramming. Statistical analysis of data was performed by mainfr"ame computer" and Statistical Analysis System pr"ogramming. 68

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69 CHAPTER 4: RESULTS OF DISSOLUTION EXPERIMENTS "That is happiness; to be dissolved into something complete and great." Willa Cather, My Antonia (also inscr-ibed on Cather's gravestone) Results of Potentiometric Study Equation 17 is the 11classic" dissolution rate expression for calcite and aragonite in seawater: ( 17) Rate This expression was utilized for the analysis of the potentiometric dissolution experiments performed at-sea during the summer of 1982. (The use of the saturation index had not been suggested during this phase of the research. However, note that the calculation of Q requires the implicit calculation of [co32-Js.) In our data analysis, K' the rate constant (k), and the spa exponent (n) were varied to find the best description of the data set. It became clear that minimum values of K'spa limited the amount of dissolution allowed by the computer model too severely, and did not allow satisfactory modeling of our dissolution experiments. The computer model predicted dissolution rates based on Equation 17 and the value of n calculated for the experimental conditions. Lower values of K' which did not lead to the pH changes which were spa

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70 observed in the actual experiments. Therefore, an examination o f the literature K'spa values was conducted Experimental values of K spa have varied significantly, and the values are related to the length of time required for equilibration of the mineral with seawater (Figure 14). When K spa is expressed in a ratio to K'spc' t he ratio varies from a maximum value of 2 .05 (Plath, 1979) to a minimum value of 1.45 (Morse et al., 1980). The variation of the K' with equilibration spa time, and our use of freshly-collected natural pteropod shells indicated enhanced dissolution rates might result from the greater solubility of a new (fresh) surface layer on the shells. Although this explanation appeared reasonable, subsequent research indicated that increased exposure time to undersaturated conditions led to enhanced dissolution rates, as will be shown subsequently The best description of the dissolution results from 84 experiments was provided by the following equation : (37) Rate (%/day) 125 (1 n)4 '1 where the value of Q was determined flith a K' /K' ratio of 2 .05. spa spc We also used a ratio of 1 78, which was appropriate to previously published kinetic studies (Morse et al., 1979; Keir, 1980) and found a reasonable description of the data with this equation: (38) Rate (%/day) 130 (1 n)3 1 Table (2) is a modified version of Table ( 1), showing how our dissolution rate predictions compare to the predictions of p r eviously

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,-.... N 'en N 0 E .._. ,... 0 x It) 10 I G I I 9 I 2 0 I I I I 8 Ill 1 8 \ \ \ 1:) 1 6 7 ....... -----------1.4 Time (days) 71 p F i gure 14 (from Byrne et. al., 1984) The solu bility behavior or aragonite obtained in vari ous investigations as a function of equilibration time. The parameter p = K spa/K1spc provides a d irect comparison of K the solubility product of calcite obtained by spc 0 Ingle et. al. (1973) (35 /00 salinity, 25 G) and K spa values (35 0 /00 salinity, 25C) obtai ne d by: 0, Plath (1979); b., Berner ( 1 976); o Macintyre ( 1 965); 0, Morse et. al. (1980). K spc values obtained in seawater at 25C and 35 /00 sali n ity agree to within 5%.

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TABLE 2 Comparison of dissolution rate equation predictions given in Table 1 compared to the equations generated by potentiometric experiments in Byrne et. al. (1984). These are Equations 37 and 38 in the text of the dissertation. Dissolution rates are exp ressed as percent of shell mass dissolved per day. [co3 2-Js = 6.469 x 105 mol/kg SW1 [co3 2Js = 7.597 x 105 mol/kg sw2 or [co3 2-Js = 8 .754 x 105 mol/kg sw3 Walter and Morse et. al. Keir Morse This work [CO 2-] 1 3 s &:J.n. 18 2 Eqn. 20 2 Eqn. 21 1 Eqn. 372 Eqn. 383 6 144 0.86 0.30 0 .05 0.77 0.87 5 .820 1.6 0. 71 0.35 1.4 1 4 5.497 2.6 1. 4 1.2 2 4 2 2 5.174 3.9 2 6 5.1 3.8 3 2 4.850 5.6 4 5 9 0 5.6 4.6 4.527 7 7 7 1 14.2 7 8 6.3 4 .204 10.3 10. 7 21.0 10. 6 8.6 3.880 13.5 15. 8 29.5 14. 1 11.3 3.557 17.3 22.4 39.7 18.4 14.7 3.233 21.6 (23.3)" 30.9 51.9 23.2 18.8 1Calculated with K = 6 .65 x 107 mol2 (kg sw)-2 For seawater spa of 35/00 salinity, 25C, and atmospheric pressure. [Ca2+J = 1.028 x 102 mol/kg sw. Units are rnmol/kg sw x 105 2Calculated with K' = 7.81 x 107 mol2 (kg sw)-2 spa 3Calculated with K = 9 0 x 107 mo12 (kg sw)-2 spa "Calculated with Eqn. 19. 72

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73 published rate equations (Morse, 1979; Keir, 1980; Walter and Morse, 1985). It is clear that the predictions agree reasonably well, (within a factor of 3 with the exception of the first line), despite differences in experimental methodology, and differing assumed K' spa values. Note especially that the predictions agree well at moderate levels of aragonite undersaturation (1 > n 0.7, where n = 0.7 a a corresponds to [co32-Js = 4.527 x 105 mrnol/kg My results, and the results of Morse et al. (1979), Keir (1980), and Walter and Morse (1985), exhibit greater dissolution rates than the results of in-situ studies. Morse et al. (1979) compared their results to in-situ estimates and found that the rates measured in the laboratory were greater by a factor of 100 than the in-situ rates. (The in-situ studies generally estimated dissolution rates by dividing weight loss by the length of exposure to undersaturated seawater.) Morse et al. states that the results from in-situ studies appear to have been inhibited by "pending" (the development of more saturated conditions within the experimental chamber) or a similar process. The length of time for in-situ studies ranged from two weeks to 250 days. Since our investigation was combined with sediment-trap sampling (Betzer et al. 1984) we applied our dissolution erode 1 to the behavior of shells settling in the ocean. An initial study of the settling velocities of representative shells consisted of observing the shells' descent through one meter in a glass cylinder filled with seawater. These observations indicated that the settling velocities of the shells could not be represented by a Stokesian settling model. Therefore, the settling rates determined for representative species

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were used in the model, and data from Vinogradov (1961) were also used The first model used a 10-layer model of a 5000-meter water column with each layer representing 500 meters of depth. The value of n and the predicted dissolution rate in each layer were then estimated. Vinogradov' s data indicated that a mid-range settling 74 -1 velocity was 1.4 em sec Combining this velocity with the predicted dissolution rates for each layer in the North Pacific Ocean provided the results shown in Figure 15. Despite the simplicity of the model, it is clear that substantial dissolution of aragonitic shells can take place in this region of the ocean. A more sophisticated model took into account the reduction in settling velocity that will accompany the loss of mass from a shell as the shell dissolves. In the case of a solid object, the size of the object would be reduced as the mass of the object was reduced. However, pteropod shells may be thought of as Caco3 "balloons", which maintain their even as the mass of the shell decreases, until the shell fragments catastrophically. Therefore, a first-order assessment of the effect of the reduction in settling velocity caused by mass loss was provided by the following equation : (39) where v is the pteropod shell settling velocity, V0 is initial pteropod settling velocity, m0 is initial pteropod shell mass, and m is the pteropod shell mass correspond ing to velocity v. When this equation is used to model the settling behavior of the shells, there is a substantial loss of smaller, slowly-settling shells in the upper

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75 2$)00 s 3,000 10 30 50 70 Of shell dinotved Figure 15 (from Byrne et. al., 1984) The cumulative% dissolution of pteropods settling at a mid-range rate, 1.4 em sec-1 at two of our station locations in the North Pacific Ocean. Significant dissolution begins at shallow depths at the northernmost station (50N, 167E). At 35N, 165E, dissolution becomes appreciable below 1000 m. As the influence of dissolution on settling rate is not considered, the cumulative dissolution shown here should be considered as a l ower -bound estimate.

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76 water column However, larger shells with rapid settling rates sink relatively unaffected to the bottom of the ocean (Figure 16a). In the northernmost Pacific Ocean, the slow settling rate of Limacina helicina shells combined with the high of lead to a of vely shallow destruction of the shells ( Figure 1 6 b) This prediction with the of shells at the top of box (Berner, 1977), and the occasional encounter with empty shells in the deep Pacific Ocean (Adelseck and 1975). The model reductions in settling velocity indicated that a shell settling with an initial velocity of 1 em sec-1 at 35N in the western Pacific Ocean would have lost approximately 30% of its initial mass when it 2200 Tne deepest traps in the study were placed at this depth, and examples of the smallest shells from this trap were selected scanning electron microscope examination This examination showed clear evidence of shell erosion on the small shells of Limacina inflata found at this depth (Byrne et al., 1984). An important observation from these micrographs is the rough, microcrystalline appearance of the eroded shell surface. This appearance has significant implications for my spectrophotometric study and also for other

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Figure 16 (from Byrne et. al. 1984) The predicted cumulative % dissolution of pteropod shells in the North Pacific versus depth for representative initial settling velocities. The initial settling rates (in centimeters per second) appropriate to each species are based on results obtained in our laboratory. Tne scales adjacent to each pteropod spec ies represent a length of 1 mm. A single pteropod spe cies (Limacina helicina) was o bserved in the highly-undersaturated waters at the northernmost station. Settling rates were determined by several m ethods (see table 6 ) Observation of the shells throughout their des cent indicated a uniform s ettling rate apparently unaffected by turbulence.

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7 8 a 1,000 36 N 1 e 5e -------b 5oN, 1111e lOOO -4,000 Of.,_.l dlsao4wec:t

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Results from Spectrophotometric Study Initial Appearance of Data Plotted Against (1-n) Following experimentation at sea with the spectrophotometric system in February-March 1985 and May-June 1985, the dissolution rate data were analyzed using the procedures discussed earlier. When the results from my experiments were first plotted against 1-n (based on 79 -7 2 -2 K'spa equal to 6.65 x 10 moles kg ), the plots exhibited anomalous scatter. In Figure 17, data from an experimental series performed on-board the R/V Marion-Dufresne is shown. When plotted against (1-n), three data points were significantly elevated above the major data distribution. When the experimental pressures were correlated with the elevated data points, there was a notable correspondence between the highest-pressure experiments and the maximum dissolution rates. Figure 17 delineates the high-and moderate-pressure regions which correspond to similar values of ( 1-Q). This pattern of maximum dissolution rates corresponding to the highest-pressure experiments was also observed in the data from other experimental series. 22Dissolution Dependence on ([co3 ]s [co3 ]) As stated before, the difference between the carbonate ion saturation co ncentration and the in-situ carbonate ion concentration, ([co32-Js [co32-]) or also expresses the degree of seawater undersaturation with respect to This undersaturation term is

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1.0 0.9 0.8 ->-0 7 <( c ....... 0.6 -w 0.5 ..... <( 0.41-a: 0 .31-0.2 1-0.1 1-0 1. 1 DISSOLUTION RATE vs. ARAGONITE SATURATION Cavollnla trldentata 45.608 mg I I I 1.0 0.9 0.8 "'?!. 1540 ,-e-: .... /ll . ,./. i 2950 I I I 0 7 0.6 0.5 Q ARAGONITE Figure 17 -Dissolution rate vs. aragonite saturation, here expressed as Qaragoni te. Experimental sequence used a singl e _g_. tridentata shell, mass 45.608 mg. Performed on-board the R/V 80 Marion-Dufresne, March 1985. This .figure shows high-and low-pressure regions (shaded areas, with pressures in psi) in the same (1-Q) range.

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similar to expressions used by other researchers to expre s s the dependence of mineral dissolution rates on undersaturation (Campbell and Nancollas, 1969; Sonderegger et al., 1976; Lasaga, 1981). Using 2-. the definitions of Q and [co3 ]s, Equat1on 36 may also be written as (40) Rate K'[CO 2-J n (1-n)n 3 s [co32-Js at pressures corresponding to the average oceanic depth is approximately a factor of two greater than its value at atmospheric pressure. If the parameters of Equation 17 and 40 are estimated from 81 experimental results at atmospheric pressure, k and K1 should have the following relationship: ( 41 ) k K' This equation indicates, for a given n, dissolution rates measured at atmospheric pressure and at a higher pressure P should have the follo wing relationship: ( 42) This relationship indicates that the predictions of rate equations based on the parameter (1-n) and those of rate equations based on the parameter t:.s will differ for variable pressure conditions. As 2-4 [co3 ]s is greater under high pressure, equation 2 indicates an enhancement of d issolution rata under pressure compared to the predicted rate at atmospheric pressure, again assuming and dentical value of n for both pressure conditions. The result o f this analysis, applied to the data used for Figure 17, is shown in Figure 18. The data distribution is shifted somewhat, so that more of the points now appear to lie on a well-defined curve.

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2-Figure 18 Data from figure 17 shown plotted against [co3 Js 2-[C03 ] (=M) "L" designates data points from the last two experiments in the sequence 82

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83 These observations provided an initial indication that is a more useful saturation index for variable pressure conditions. However, four points (labeled L) remain elevated above the curve. These points indicate the presence of an important process which will be discussed in a subsequent section. One reason for the improved appearance of the curve is the contrasting behavior of (1-n) and in response to pressure. To demonstrate this behavior, a hypothetica l sample of seawater which was exactly saturated at atmospheric pressure (at 5C and 35 o I 0 0 salinity) was taken to successively higher pressures using our calculation algorithm. All conditions other than pressure were held constant. Figure 19 shows that increases at an increasing rate with higher pressures, while the rate of increase for (1-n) decreases as the pressure is elevated. As the undersaturation of oceanic waters increases markedly with increasing pressure, and aragonite dissolution rates are directly related to the increasing degree of undersaturation, appears to better express the change in the saturation state with pressure, and therefore should show an improved correlation with more rapid dissolution rates observed under pressure. All of the datasets generated in this research were analyzed by the least-squares method to compare the differing predictions of (1-n) and as dissolution rate indices. All of the data were fit to an equation of the form R = Axn (Eqn. 35) In each analysis method, data sets were in the form of x,R data pairs, with x given the value of either (1-n ) or and R given the value of the corresponding dissolution rate. The observed dissolution rates, R, were identical for both sets of data.

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84 1 1 1 1 X CY X 10 10...1. I 0 -"-,. (> 9 Q Cl? X 0 ..... 0 >< X (]1 8 8 a_ 0 X " 'It CJ) Figure 19 6(1-n)/6P and 6(6s)/6P as calculated for a hypothetical sample of seawater (35 /00 salinity, 5C, TA = 2.400, rco2 = 2 3387), which is exactly saturated at atmospheric pressure. 6(6s)/6P increases with increasing pressure, in contrast to 6(1-n)/6P, which decreases with increasing pressure.

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85 The results of this analysis for fifteen dissolution rate datasets are shown in Table 3. In 13 out of 15 cases, the residual sum-of-squares value (RSS) for the datasets is lower than the RSS value for the (1-Q) datasets. This result shows that the relationship between the seawater saturation state and aragonite dissolution rate is best expressed using the saturation index when the rates are determined under variable-pressure conditions. Evidence of Surface Effects Compli cation of Functional Dependence In several of the dissolution rate curves that have been generated in this research, data points which appeared to lie substantially above the "best" rate curv e were generally found to be from expe riments performed late in an experimental sequence. The four points labeled "L" in figure 18 are representative examples Elevated dissolution rates near the end of a long experimental series were the first indication of a process whic h affects the dissolution rate of a shell as erosion of the shell surface progresses The effect becomes more evident the longer a shell is exposed to undersaturated conditions. With the ex ception of sequences where a fairly large amount of aragonite was used f o r a relatively sh ort period of time, this effect influences the quality of the data that is used to generate the analyzed rate curves. This effect is also distinct from "breaching", which is the opening of chambers in objects (notably foraminiferal tests) which will increase the surface area of a

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TABLE 3 -Values of n and residual sum-of-squares (determined by least-squares analysis) for dissolution rate vs. 6s datasets, AND n and residual sum-of-squares for dissolution rate vs (1-Q) datasets. 6V = -31.3 cm3/mol. (Subsequent analyses employed a different value of 6V, the molar volume change for aragonite dissolution.) ( 1-Q) Dataset n RSS RSS n 1 141 .35 0.2265 0 .493 1. 448 72 2 1 .258 13 0.0426 0.064 2.116 .30 3 0.96 .27 0 .278 0.487 1 .056 .42 4 0. 91 41 1.509 1. 342 1.37 .55 5 1. 728 .25 0.0772 0.1026 2.163 .38 6 1. 379 31 2.044 2.69 2 .678 .72 7 1. 496 16 0.6 0.396 3 .561 .33 8 1.616 18 1.27 5.83 2.51 .675 9 2.10 .75 1 .066 2 .066 3 .686 2.10 10 0.964 .22 0.281 0 .378 1. 447 41 11 1.673 26 0 .042 0.061 3 .25 .65 12 0.84 .55 6 .65 7.83 0.66 .55 13 2 .13 .53 3.51 9.95 1 .205 .75 14 1.29 15 o. 1015 0.1668 2 .075 .33 15 1.706 .37 2.225 6.993 0 .897 .49 86

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particulate assemblage. In our experiments, the use of simple (Cavolinia tridentata, Cavolinia gibbosa Cuvierina columnella, Clio pyramidata) eliminated the influence of "breaching". Progressive Exponential Increase in Long-Term Experimental Series In order to clearly delineate the effects of dissolution on the dissolution rate of an aragonitic shell, a series of 10 experiments was performed over 9 days. A single assemblage of a Cavolinia 87 tridentata and Cavolinia gibbosa shell was used in all 10 experiments. Each experiment consisted of three rate determinations at three suc cessively higher pressures, so that each experiment generated three data points. The ten experiments were analyzed in three groups; the first three days, the second two days (3 experiments), and the last four days. The raw data from these experiments is shown in Figures 20, 21, and 22. In the following figUJ:'e (Figure 23), the rate curves generated by least-squares cUJ:'ve fitting of the data are shown. During this sequence the shells were stored in the same seawater sample that was used in the experiment. Thus, the shells were constantly exposed to undersaturated seawater. In these experiments, the dissolution rate curves exhibited a clearly increasing trend with time. The entire length of exposUJ:'e was approximately 200 hours. In a second 9-experiment sequence over 9 days, the least-squares fitted curves did not increase as dramatically with time as in the above sequence. However, the best-fit cUJ:'ves for the data from this sequence still indicated increasing dissolution rates as time progressed. Figure 24 shows the curves for the three data groups in

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88 Figure 20 Data from the first three days of experimentation of a 9-day, 10-experiment sequence with one assemblage of a f. tridentata and C. gibbosa shell.

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4 Increasing rate + /"\. curves with >3 extended time a: 0 + + v 2 UJ 1-a: 0::: 1 .. + + I I 10 20 30 40 80 DELTA S Figure 21 Data from the next two days (three experiments) of the 9-day sequence using the same shell asse m blage 89

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5 Increasing rate ,..... curves with >-time + extended a: c .._, a: ... ... 1 + t I I I 10 20 30 40 50 60 DELTA S Figure 22 Data from the final four days of 9-day sequence using the same shell assemblage. 90

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,-.. >-3 a: 0 v2 LLJ 1-a:: et::l Increasing rate curves with extended time 10 20 30 DELTA S 3 1 40 50 60 91 Figure 23 -Increasing rate curves from 9-day 1 a-experiment sequence Curves were generated by least-squares analysis of data from Figures 20 (1) 21 (2) and 22 (3)

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"' >-a: 0 v L1J t-a: 92 1. 5 1 5 10 20 30 40 50 60 DELTA S Figure 24 -Increasing rate curves from second 9-day t 9-experiment sequence. t first three days; +t second three days; ot final 3 days This sequence also shows increasing dissolution rates with extended exposure to undersaturated seawater.

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93 this sequence. Experimental procedures should have limited the influence of surface effects on the majority o f the data. When uniform solution conditions comparable to oceanic levels of under-saturation are maintained, the surface effects which cause increased dissolution rates develop slowly, as demonstrated by the curves in Figure 23 and 24. In the two experimental sequences in which increasing rates were observed, the shells were exposed to undersaturated seawater for about 200 hours. For most of the data collected in this research, the total exposure time was approximately 40-50 hours. In the ocean, rapidly settling shells will reach the ocean floor in 2 to 3 days. However, shells with slow settling rates will be exposed for longer periods, which will allow the development of surface alterations during settling. Continuous experimentation at-sea and passive laboratory storage in seawater also limited such effects by minimizing the length of "active" dissolution periods. "Active" dissolution means dissolution in a well-mixed sample chamber. "Passive" refers to storage of a shell in an unmixed container, where local saturation near the shell surface could inhibit dissolution. Sur face effects apparently are present in all of the experimental sequences shown in Table 3. Examination of the two long-term experimental series illustrated in Figures 20-24 indicates the maximum magnitude of influence on the experimental results from these effects. Because of the extended duration of exposure to undersaturated seawater, the influence of changing surface chemistry was largest in these experiments. Therefore, the effects demonstrated in Figures

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94 20-24 should be larger than the surface influence in the majority of the experimental sequences. Datasets 1-3 in Table 3 were used to generate Figure 24. Datasets 6-8 correspond to figures 20-22 and the curves shown in Figure 23. For five of the six datasets from the long-term experimental sequences described above, the RSS value for the 6s dataset is less than for the (1Q ) dataset. This analysis clearly shows, that even in the presence of surface processes which appear to accelerate dissolution rates, the model based on 6s best describes the dissolution behavior observed in the experiments. (Since the surface changes influence dissolution rate, the effects are a constant factor in the data regardless of which saturation inde x is chosen for the analysis.) It is believed that surface alteration is the most likely explanation for the observed variation in dissolution rates. The rate constant k varies with pressure (see Equation 40), and the exponent n is empirically derive d from the data analysis method. It would be instructive to examine the dependence of dissolution rates on 6s at atmospheric pressure by varying the solution composition, to determine the validity of 6s under all conditions. Thus, even considering surface effects, the saturation index 6s best expresses the relationship between aragonite dissolution rates and the seawater saturation state for condi t ions of elevated, variable pressure. The use of 6s under these conditions is a fundamental result of this research. In a following chapter, the manner in which surface effects influence the observed dissolution kinetics will be further described.

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Influence of t:.V Analysis of Dissolution Rate Data An aspect of the data analysis is the value of t:.V dissolution at 2-5C which is used in Equation 28 to calculate the of K' with The data in table 3 spa were with t:.V = -31. 3 cm31mol (Mucci et al., 1982; Feely et 95 al., 1984). The maximum value of t:.V is the value, -39.5 . cm3 /mol. This value is found by the difference between the and calcite molal volumes, 2 8 cm3/mol 1979), the t:.V for calcite dissolution, -42.3 cm3 /mol (Ingle, 1975; 1979). The next section examines the influence of t:.V on the data analysis. Equation 28 is shown again clarity in the follow ing discussion: (28) ln (K p / K 0 ) spa spa -(t:.V/RT)P Results from columnella Initially, a value of t:.V dissolution was not a goal of this examination of a high quality dataset obtained in this allowed a new of this t:.V value The dataset was obtained at-sea the R/V in June 1985. columnella shells used in an sequence which was completed in two days was continuous, for a total exposure time of 45 hours. (Only one dataset in table 3 has a

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total exposure time, but only six data points were obtained in the shorter interval.) Examination of the influence of 11V on the data r-evealed that larger values of /1V (up to -40.0 cm3/ ool) led to a tighter distr-ibution of data points about an appar-ent best-fit curve. The visually apparent impr-ovement was confir-med by least-squares analysis of the data generated by different values of 11V. 96 The residual sum-of-squares values whic h were generated by least-squares analysis of the data, were examined as a function of /1V. When the dissolution r-ates were expressed as a function of 11s, the residual sum-of-squares achieved a minimum value. However, when the data was expressed as a function of (1-n), no minimum RSS value was found in the range of reasonable values for /1V. Figure 25 shows the contrasting behaviorof the residual sum-of-squar-es analysis for the two saturation indices. The minimum value of the residual sum-of-squares was found when /1V was equal to -37.0 cm3 /mol. Figur-e 26 shows the dissolution rate data plotted against 11s, generated with our best-fit value of /1V. Note that this value lies approximately midway between the highest value of /1V exper-imentally determined in solubility studies, -33.1 cm31 mol, and the highest "theor-etical" value of /1V, -39 5 c m31 mol. I believe that the value of /1V generated by this analysis indicates that previous experimental determinations of /1V have been low. The actual reason for the discrepancy between the exper-imental values and the "theoretical" value is unknown. This discrepancy may be related to the difficulty of distinguishing, for aragonite, between an appar-ent equilibrium (or steady-state condition) and a true, thermodynamic

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97 6 0 5.0 (/') LU 0:: 4 0 <( => 0 (/') LJ.. 3 0 0 => (/') 1 0 ..J <( => R= Q (/') .8 LU 0:: .6 .4 -34. 0 -36.0 -38. 0 -40. 0 -42. 0 -44. 0 ll.V ARAGONITE (cm3 /mol) Figur-e 25 Var-iable pr-essur-e ar-agonite dissolution r-ate data fr-om Cuvier-ina columnella exper-iments is examined using Eqns. (17) and (35), and Eqns. (36) and (35). Foreach model, the r-esidual sum of squar-es for ourfitted data is shown as a function of 6V.

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98 RAT E (%/DAY) 1\) 0 co ..... ""' ,.., () 0) 0 (t.) 1\) 1\) t O L...l (/J I ,.., () 0 co (t.)f\) 'Ct.> L-11\) 26 -Dissolution data C. columnella shown as a function of [co32-Js -[co32-J 6s), using best-fit 6V estimate of -37.0 cm3/mole.

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99 state of equilibrium. Only a few other datasets exhibited a sufficient range of saturation values and pressures to qualify for this type of analysis. However, a factor which will determine the success or failure of this type of analysis is the pressure range of the experimental sequence. Changing the value of causes shifts in the position of the points relative to the horizontal or 1-n) axis, and the dissolution rates do not change Maximum shifts in the points only occur for experiments performed at very high pressures, and experiments performed below 1000 p.s. i. have negligible shifts. In order to achieve a minimum in the residual sum-of-squares as a function of there must be enough points with significant position shifts as is changed to influence the analysis. Furthermore, any other factors that affect the dissolution rate make the analysis more difficult. Experiments at pressures higher than those acheivable with this system would also help to fix the value of Due primarily to the above factors, no other datasets were found that yielded a minimum RSS value when analyzed according to the method used for the Cuvierina columnella data. The analysis of the Cuvierina columnella data, which shows an improvement of the apparent relationship between dissolution rate and with a higher value of supports the higher "theoretical" values of the molar volume change for aragonite dissolution. Careful experimental design specifically for this topic could yield more precise results. To investigate whether higher values of also improve the data in the same manner as is demonstrated in Table 3, all of the datasets were recalculated using the maximum theoretical value of -39.5

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100 cm3/mol. (The theoretical value was obtained by adding the difference between aragonite and calcite molal volumes, 2.8 cm3/mol, to the value of for calcite dissolution, -42.3 cm3/mol.) Each new ( 1-n) and dataset was then submitted to least-squares analysis to generate new RSS values. These values are shown in Table 4. Table 4 indicates that the decrease in the RSS values for 6s datasets as compared to (1-n) datasets is also observed for < -36.5 cm3/mol. A compariso n of the value s in Tables 3 and 4 indicates two other noteworthy points. The first is the reduced values of the RSS values in Table 4 compared to Table 3. For the datasets, in ten out of fifteen cases the RSS value in Table 4 is less than in Table 3, indicating a better curve fit for the higher value of For the ( 1-n) datasets, thirteen out of fifteen datasets exhibit an improvement with the higher value. The second point is the value of the exponent (n) The majority of the datasets in Table 4 show r ed u ced values of the e xponent compared t o the data in Table 3 These exponential values are disc ussed below. This type of analysis also indicates that kinetic measurements under pressure can provide insight into fundamental physico-chemical constants. To further illustrate the trends in the data, the "best-fit" value of (-37 0 cm3/mol) was also used to calculate the saturation parameters and (1-n). The RSS values for each dataset for the three values of are shown in Table 5 Table 5 demonstrates that higher values of improve the functional relationship for the majority of the datasets. Variation in the data quality and the

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101 Table 4 Values of n and residual sumof-squares (determ ined by least-squares analysis) for dissolution rate vs. 6s datasets, AND n and residual sum-of-squares for dissolution rate vs. (1-n) datasets. 6V = -39.5 cm31mol. 6s ( 1-Q) Dataset n RSS RSS n --1 .083 .27 0.140 0 .330 1 .308 .51 2 1 .077 12 0 .047 0.061 1. 891 .27 3 0.995 .27 0.184 0 .371 1 .165 .49 4 0 .790 36 1 .525 1.555 1.030 .50 5 1 .559 .29 0.124 0.076 2.361 .35 6 1 .184 .26 1 .958 2.316 2.426 .60 7 1.236 14 0.548 0.481 2.884 .29 8 1. 351 12 0.909 4.146 2 .380 .52 9 3. 142 .99 0 .659 1. 761 4.580 .63 10 0.330 19 0.282 0.36 1 1 392 38 11 1 .442 .21 0 .037 0 .055 3.039 .58 12 1 .016 .57 6.022 7.389 0 .806 .60 13 1 779 .29 3.213 7 .831 1 .651 .80 14 1 135 12 0 .090 0.144 1. 930 .30 15 1 .624 15 0.526 3.938 1. 764 .59

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TABLE 5 Residual sum-of-squares values for non-linear least-squares curve fits for rate vs. b.s datasets 1-15 (shown in tables 3 and 4). Data was generated for three values of b.V; -31.3, -37.0, and -39. 5 cm3/mol Dataset b.V = -31. 3 b.V = -37.0 11V = -39.5 1 .2265 1744 140 2 .0426 .04522 .047 3 .278 .2088 .184 4 1.509 1.5153 1 .525 5 .0772 .10209 .124 6 2.044 1. 9843 1. 958 7 .6 .62768 .648 8 1.27 1 .0418 .909 9 1 .066 794 .659 10 .2814 .2824 .28186 11 .042 .0386 .037 12 6 .65 6.205 6 .022 13 3.51 2.3496 1 .883 14 1015 13116 .090 15 2.225 .38846 .526 ** -designates decreasing RSS values with higher b.V. ** -designates minimum RSS value at b.V = 102

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103 number of high-pressure points (which are most affected by changes in the value of may be the reason that the improving RSS trend is not observed in all of the datasets. Range of Exponential Values Previous research has produced a range of exponential values in rate equations employing (1-n) as the saturation index (Eqn. 17). The maximum value found for aragonite (Keir, 1980) is 4.1. The minimum value of the exponent for aragonite in previous laboratory experiments is 2.5 from Morse and Walter ( 1985). When my dissolution rate data were analyzed using as the saturation index, and was equal to the -31.3 c m 3/mol, two data sets yielded a value of n approximately equal to 2. 1 (2.10, 2.13). The majority of n values ranged between 1 and 2. The lowest value of n determined was 0 84, and three other data sets gave values of n slightly less than 1 (0.91 and 0.96 for two sets). The average value of n for the fifteen datasets analyzed (Table 3) was 1.41. For the data in Table 4, which was generated with equal to -39.5 cm3/mol, the average value of the exponent drops slightly to 1. 36. These values of the exponent are significant to model predictions of the dissolution process in the ocean, as well as the conclusions which will be stated at the end of Chapter 5. Dye Concentration Experiments The spec trophotometric method of determining seawater pH and the

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104 change in seawater pH caused by aragonite dissolution requires the use of relatively high concentrations of the pH-sensitive dye phenol red The concentration used in these experiments was approximately 8 x 10-6 molar A concern involving the use of this new method was the possible effect of the dye on the experimenta l results. To test the effect of the dye's contribution of H+ to the system, a term for the concentration of the dye's protonated and deprotonated forms as a function of pH was added to the computer program used for data analysis. This term allowed calculation of the effect of the dye on the pH of the system The strongest effect from the dye was found when the pH of the system was equal to the pK1 of the However, this effect affected the calculated pH change in dissolution experiments by less than 0 .1%. In order to further test the possible effect of the dye concentration on the results obtained in this research, a short group of experiments was devised with varying dye concentrati ons. These experiments were designed to investigate the possibility that the dye could adsorb to the shell surface, and cause a misinterpretation of the dissolution rate determinations. The saturation state in the experiments was kept as constant as possible. Three concentrations were used; the normal concentration used in the majority of the experiments, and concentrations that were one-half and twice the normal concentration 'rlhen the dye concentration was twice the normal concentration, the absorban ce at 558 nanometers was too high to be measured In this case, the absorbance was measured at 575 nanometers, where the absorbance was lower. The equations used to calculate the pH remain the same, as

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105 long as Amax is measured at the same wavelength. Figure 27 shows the results of this series of experiments. No effect of the varying dye concentration is evident from this series. The method used to calculate the error bars in Figure 27 is described below. Error Analysis Precision of Dissolution Rate Determination In Figure 27, which shows the results from the dye concentration experiments, estimated error bars are also shown. The calculation of the dissolution rate was made by calculating the pH change from the absorbance data, calculating the change in the seawater Ico2 concentration from the pH change, dividing this value by the length of the experiment, and then using the mass of the shell to determine the dissolution rate in percent per day. In order to assess the possible error in this measurement, the uncertainty in the absorbance measurements was estimated as .001 absorbance units. If the initial absorbance value is assumed to be .001 absorbance unit lower than measured, and the final value is assumed to be .001 absorbance unit higher, the maximum absorbance change could be .002 absorbance units more than the measured change. Conversely, the minimum absorbance change could be .002 absorbance units less than the measured change. The data was adjusted to reflect these differences, and the computer program then calculated maximum and minimum dissolution rates. These

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z4 t-4 :E: ..... w .... :E: .... 28 32 36 DELTA S 106 Fi gure 2 7 -Results from experiments in which the phenol red concentration was varied between 1/2 and 2x the normal experimenta l concentration. 1/2 -one-half normal co ncentration; N norma l concentration (8 x 10-6M); 2x-twice normal concentration.

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107 error estimates are proportionately greater when dissolution rates are slow, which is an expected result, as it is more difficult to determine small absorbance changes. Precision of Alkalinity Determinations The determinations of Ico2 and total alkalinity performed at sea are described in Feely et al. (1984). The maximum uncertainty of the 2co2 determination is 1 2 llWkg, and for total alkalinity the uncertainty is 5 lleq/kg These uncertainties lead to a maximum uncertainty in Q for aragonite of 5 percent. Since the quantities used to calculate are the same, the estimated uncertainty for is also 5 percent. In the laboratory, a titration method (Culberson et al., 1970) was used to determine the alkalinity, and the pH of the seawater was determined potentiometrically. 2co2 was calculated from the carbonate alkalinity and pH. The uncertainty of the alkalinity titration is estimated to be 8 lleq/1. The uncertainty i n the measurement of the seawater pH is .01 pH units. The uncertainty in the alkalinity is more important than the uncertainty in the pH for the calculation of The calculated Ico2 also has an uncertainty of 8 lleq/1. These uncertainties give approximate 6 percent uncertainty in the value of Q and For example, Q = 0 .6 03, and = 25 1 .25.

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108 CHAPTER 5: EFFECT OF SURFACE ALTERATION ON DISSOLUTION RATES "The secondar-y imagination . dissolves, diffuses, dissipates, in or-derto r-ecr-eate; or-wher-e this pr-ocess is r-ender-ed impossible, yet still at all events it str-uggles to idealize and to unify. It is essentially vital, even as all objects (as objects) ar-e essentially fixed and dead." Samuel TaylorColer-idge, Biographia Liter-ar-ia . n Held Constant at Low and High Pressur-e In an effor-t to demonstr-ate that 6s is a super-iorsatur-ation index to use for-var-iable pr-essure r-ate deter-minations (and for modelling dissolution r-ates undervariable pr-essure), two sets of exper-iments wer-e designed to pr-ovide dir-ect compar-ison of the applicability of 6s and (1-n) for-var-iable-pr-essure conditions. In the fir-st set of exper-iments (r-efer-r-ed to as "mano-a-mano" exper-iments), pair-s of seawatersamples ("A and B") wer-e pr-epar-ed with measur-ed additions of HCl. Ideally, in these exper-iments the value of n at atmospher-ic pr-essure in sample "A" would be equal to the value of nat elevated pr-essure in sample "B". Eac h pairof exper-iments was per-for-med on the same day, with the atmospher-ic pr-essure exper-iment pr-eceding the pr-essurized exper-iment. The pressure in the second exper-iment was adjusted so that the values of n in the two exper-iments would be as close as possible. Additionally, in two multi-pr-essure experiments the satur-ation

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109 state was nearly equal at two different pressures (due to aragonite dissolution at the lower pressure). These results were added to the "mano-a-mano" experimental results. When the six pairs of results were expressed as a function of (1-n), the dissolution rate at elevated pressure was greater than the dissolution rate at atmospheric pressure, for all six cases. When the experimental results were expressed as a function of the faster dissolution rate of each experimental pair correlated with a higher value of Held Constant at Low and High Pressure In another set of experiments designed to compare the applicability of (1-n) for variable pressure conditions, pairs of seawater samples were again prepared with measured additions of HCl. In this case, the samples were prepared so that the value of for the low-pressure solution and the high-pressure solution would be approximately equal. (These experiments were referred to as "reverse mano-a-mano" experiments). Since both of the experiments were performed at a pressure higher than atmospheric pressure, the order in which the experiments were performed was varied. In these experiments, our theory predicts that dissolution rates for each pair should be the same. This behavior was observed in two of the five experiments. In one case, the rates were virtually identical, and in the other case, the rates agreed within experimental error. However, the other three experiments showed different rates. In each of these cases, the experiment performed second, whether at

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low high had a slightly dissolution The maximum in dissolution in the second was 25 of the in the When as a function of ( 1-n) 1 these conflicting. Two lent to the (1-n) model, with dissolution to values of ( 1-n). Three conflicted with based on (1-n). 110 1 the of this set of did not conclusively that 6s was the saturation index conditions. The of the n was held constant followed the expected with dissolution the high but the in which 6s was held constant yielded mixed Although these lend some to the usefulness of 6s, they also indicate that dissolution are affected by slow chemical The statistical data in Tables 3 and 4 evidence the 6s model (Eqn. 36). the possibility of changes in dissolution caused by of the shell complicates the analysis of the mte data as a function of the saturation indices 6s and (1-n). This could be the nature of the data in a long of The in with time, as manifested in the mano-a-mano" can the exponential curvature of the data This was consistently in my

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111 Effect of in Solutions In experiments at sea, was continuous, and a shell set of shells) was used in experiments and then breaks between common during To the oceanic dissolution the shells in undersaturated In the initial stages of the a solution was made highly by. the addition of HCl. The amount of HCl added to the solution in which the shells was twice the amount to oceanic levels of The of in this solution was the highest to which the shells exposed at any time the When a five-day sequence of multiple-pressure at successively was using this between unusual obtained 28). The observed dissolution the sequence, and elevated the same level of on the final days as compared to the days. As this sequence it was also noted that the change in the twenty minutes of a was than the change the of the in the last the in the conducted at the lowest exceeded the at the next highest

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(!) lSI 'P"4 x3 r-. z t---4 :E: ..... 132 L1J E .. _,. UJ f-1 a: 20 28 36 44 52 DELTA S 60 112 68 76 Figure 28 Experimental sequence during which shells were stored in highly-undersaturated seawater solution between experiments, showing strong effects of advancing dissolution. o Day 1 ; 0 Day 2 ; +, Day 3; X, Day 4; o, Day 5

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113 Because of the high degree of undersaturation in the seawater storage solution, which may have caused extensive alteration of the shell surface, the effects observed here are exaggerated beyond the effects which would be expected from the natural dissolution process. The wide variation in dissolution rates in this dataset made it impossible to analyze by the least-squares curve-fitting method. The unusual effects observed in this sequence may be similar to some of the results obtained when the 6s dependence of the dissolution rates was dicectly tested ( "cevecse mano-a-mano" experiments). As stated befoce, in thcee of the five dicect compacison expeciments, the dissolution cate measured in the second expeciment was greater than the cate detecmined in the first experiment. In Figuce 28, note that the Day 1 points define a reasonable cucve, with incceasing dissolution rates cocresponding to an inccease in the degree of undecsaturation. This cucve is similar to the cucves shown in Figuce 23. As sucface altecation pcogressed, drastic changes in the relationship between dissolution cate and saturation state became evident, as shown by the data from days 2-5 in Figure 28. Thus, Figuce 28 pcovides a clear demonstration of the effects of surface alteration on dissolution rate. In contrast, Figure 23 shows how milder alterations (similar to natural changes) influence the rate vs. saturation state data. Based on this observed behavior, it is clear that the experiments described in this dissertation investigated the first-ordec influence of seawater saturation state on dissolution rate, with the second-order effects of surface alteration integcated over the length of each experimental sequence. In some cases, these effects were

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114 minimal, and in other cases, they were significant. Obviously, when these effects became significant, the process caused an increasing degree of scatter in the data. Due to the scatter induced by these effects, the only way to clearly distinguish between the predictions of the (1-Q) indices is through a statistical comparison, whic h has been shown previously. Physical Nature of Shell Surface Alterations SEM examination of shells taken from deep sediment traps revealed the. surface characteristics of shells which had undergone moderate dissolution. The initial smooth surface layer of the shell was removed, and the underlying layer was rougher, appearing to consist of microscopic aragonitic crystals approximately 1-5 in length. Other researchers have reported that increasing surface area led to an apparent increase in dissolution rate (Morse et al., 1979). Thus, the increasing trend of the dissolution rate curves should be expected as the exposed surface area of the shell increases. Furthermore, dissolution of the roughened surface provides an increasing number of high energy sites on the crystal surface where ion detachment will be favored. These sites are termed "steps" or "kinks" by Berner and Morse (1974). Morse (1986) states that the solubility of calcite increases exponentially with decreasing particle size. This effect is not very significant until the particles are quite small (less than 0.1 diameter, Figure 29). However, the presence of such small particles will cause a noticeable increase in the measured solubility of the

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t 7 .. "' 0 ... IC .. ... u ... .. u M ... BULK SOLIO SOLUIILITY l:DG! LMGTH (uM) Figure 29 (from Morse, 1986) Effect of decreasing grain size on t h e the bulk solid solubility of ca l cite. 115

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116 mineral (Sass et al., 1983; Walter and Morse, 1984). The solubility increase is due to an increase in the free energy of a mineral, which is a function of the molar surface area S (Stumm and Morgan, 1981). Therefore, the development of eroded, microcrystalline areas on an aragonite shell could alter the solubility of the shell. Figure 30 shows two shells which have been affected by dissolution. Figure 30a is a Limacina inflata shell taken from a sediment trap deployed at 2170 meters in the North Pacific Ocean ( 35N, 165E). The shell's roughened surface demonstrates the increased surface area and small crystal size of the exposed surface. Also note the central portion of the shell. The Limacina family has two different crystal structures composing the shell. A smooth layer of "prismatic" crystals is underlain by "crossed-lamellar" crystals (Be' and Gilmer, 1977). The remaining portion of the prismatic layer appears at the center of the Limacina shell. (Other views of Limacina inflata shells taken from sediment traps appear in Byrne et al. 1984). Figures 30b-c are of a Cavolinia tridentata shell which was used in potentiometric dissolution experiments during the 1982 cruise. This shell was used in two long experiments, and thus showed only initial stages of dissolution. Figure 30b shows the intersection of two ridges on the shell where dissolution has begun. Figure 30b is a closeup of this region, showing the exposed ends of the aragonitic rods which are the structural units of the shell. Be' and Gilmer ( 1977) also show the microstructure of Cuvierina columnella and Cavolinia longirostris shells, which have a helical structure of intertwined aragonitic rods.

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Figure 30 (top) Limacina inflata shell taken from 2170 meters in the North Pacific Ocean, showing the effects of dissolution on the surface and the roughened nature of the partially-dissolved surface. (middle) Ridges on a Cavolinia tridentata shell used in dissolution experiments, demonstrating initial dissolution effects. (bottom) Closeup of ridge surfaces, showing roughened surface characteristic of exposed aragonite rods

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118

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119 Thus, the roughened, inhomogenous areas on a shell produced by dissolution may caus e changes in the solubility of the aragonitic surface in two ways; one, by increasing the surface area of the shell, and two, through an increase in the solubility of the mineral, similar to the effects of decreasing particle size for a mineral powder. It should be fruitful to investigate the effects of an increasing K' spa on the analysis of data from some long-term experimental sequences. Effect of Increased Values of K' on the Data Analysis ------"---'As previously noted, in some long experimental sequences the dissolution rates from experiments performed later in the sequence were elevated over the data curve which included the majority of the data points. An example of this behavior has been shown in Figure 18. In order to estimate the magnitude of the change in K'spa which could cause the observed changes in the dissolution rate data, the points from the "later" experiments (labeled "L" in Fig. 18) were analyzed separately from the points in the primary data distribution. This method of analysis is illustrated in Figure 31, using the same dataset shown in Figures 17 and 18. In Figure 31a, the data from the first days of experimentation follows a well-defined curve, but the four points which were generated on the last two days of the experimental sequence (+) clearly exhibit an increased dissolution rate compared to the rest of the curve. When these points are analyzed with an increased value of K' they are translated to the right, as shown spa in Figure 31b. By increasing the solubility product, all of the data points appear to exhibit approximately the same rate dependence on 6s.

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121 In Figures 32 and 33, the results of the same analysis method are shown for other data sets which exhibited similar alteration of dissolution rates with time. In each case, it is clear that a larger value of K'spa' implying greater solubility of the pteropod shell surface, improves the r-ate curve and in many cases r-educes the discr-epancy between the initially determined r-ate curve and the laterdata points. K'spa determined by Mucci (1983) is 6.65 x 107 mo12 kg2 The largest K'spa requir-ed to align the data points to the rate curve is 8.05 x 107 for a maximum K' /K' r-atio of 1.75. This estimate spa spc assumes that increasing surface area does not contribute to the increased dissolution rates, and ther-efor-e must be considered an upper-bound estimate. In the other thr-ee analyzed cases, K' falls spa between 7.36 and 7.82 (x 107 ) mo12 kg2 These results ar-e quite interesting in light of previous determinations of the aragonite solubility product in seawater. Morse et al. (1980) correlated all of the K' spa determinations which pre-dated their investigation, and showed that the value of K'spa decreased with longer equilibration times. The minimum K' of 6.65 spa x 1 o-7 mo12 kg -2 was found in their research after a period of equilibration in excess of 80 days. The K' IK' ratio, compared spa spc to the K'spc in seawater determined by Ingle (1973), varied from a maximum of 2.05 to the minimum of 1.525. (Figure 14) It therefore appears that the value of K' decreases with spa longer equilibration times, and the value of K'spa increases with longer dis-equilibration times (i.e. exposure to undersaturated seawater). In Berner and Morse's (1974) theory of calcite

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122 "" e >2 a: a ....... e v 0 0 00 LLI 11 0 a: ie 0 0 + 0 00 8 16 2 4 3 2 40 48 se DELTA s 3 + "" e >2 a: a e v 0 e ee LLI 11 0 a: e e + 0 ee 8 16 2 4 32 40 48 56 DELTA s Figur-e 32 -a. Data fr-om initial exper-iments using a C tr-identata shell, mass 1 2.50 2 m g (+) Last two days o f exper-iments. b Data fr-om last two days adjusted with incr-eased value o f K spa

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123 5 4 +a /'\ + >a: 0 3 t a .t + v 2 e (l) w e .,_ a: .. 0::: e 1 .+, e 0 ee 8 16 24 3 2 40 48 56 64 72 DELTA S C' ..; 4 i /'\ >- a: 0 3 t e -I + v e (l) w 2 e .,_ .. a: 0::: 1 -i+ + e 0 e e 8 16 2 4 3 2 48 48 56 64 72 DELTA S Figure 33 Data from 9-day 10-experiment sequence, divided into thr ee groups correspondi n g t o F i gures 20, 2 1, and 22. 0 first three days ; + next two days; + last 4 days b Data adjusted with increased va l ues of K spa

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124 dissolution, the active sites for ion detachment from the mineral surface were believed to be inhomogeneities on the surface, which were called "kinks" or "steps". The equilibration process probably reduces the number of such active surface sites, in effect "SIOOothing" the surface to a minimum solubility. The process of dissolution appears to act in reverse, increasing the number of active surface sites by eroding and roughening the surface. In modeling the results of our potentiometric investigations of dissolution, we invoked the idea of higher solubility products for the aragonite shells due to their "freshness" or "naturaL11ess", as they were taken directly from the ocean in sediment traps. The results of subsequent investigations now appear to indicate that the higher K'spa was a valid idea, but the reasons for the higher value are different. During the course of the potentiometric experimentation, a small number of shells was used in several experiments, with almost continuous exposure to undersaturated water under high-pressure conditions. In fact, seawater from near the depth of maximum undersaturation was usually used in the experiments, but the pressure was substantially increased over in-situ conditions, with the resulting levels of undersaturation highly elevated. The process of surface activation described above, which appears to increase the value of K' and correspondingly enhances the dissolution rate, spa should have been operating to an equal or greater extent in the potentiometric experiments. Visual examination of the shells used in the potentiometric experiments shows an almost complete opacification of the shells, which are nearly transparent in their natural state. The development of nearly opaque shells, also called "etching", does

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not occur when the shells in when the shells with the in the of has indicated that opacification accompanies lengthy exposure times to conditions. 125 Therefore, the in dissolution time which has been in this appear's to be the of two The first is an in the surface of the shell due to caused by the dissolution and the second, slower is a in the value of K' spa wit!:). lengthening exposure times to These processes are in with pt'evious investigations of aragonite solubility, the physical of the shells, and the molecular-level of dissolution. using the 6s conditions only could also indicate surface is the only which dissolution or if is an additional influence of pt'essut'e. Influences of Surface It is the delineated in this that dissolution of the shell surface can influence the dissolution kinetics of a shell in single as well as in sequences. In both cases, the net effect is to the dissolution In the which provided the of the data, the effects of

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126 alteration were most evident at highest pressures and the correspondingly highest degrees of undersaturation. Furthermore, increased dissolution rates became more apparent in experiments which occurred at the end of a long-term sequence. In summary, I propose that in the complete absence of changing surface chemistry, aragonite dissolution rates would exhibit a linear dependence on the quantity 6s which expresses the degree of seawater undersaturation. This inference is supported by the initial experiments in long-term sequences, which resulted in rate equations with exponential values near 1. (Datasets 1, 6, 12, and 14 in Tables 3-5 are groups of such initial experiments.) I t is possible to envision an idealized experiment, in which an instantaneous determination of aragonite dissolution rates is made over a continuous range of undersaturation values, so that surface chemical changes will not occur In su ch an experiment, dissolution rates are determined for every point over the 6s range in brief (e.g. millisecond) experiment. The results which I would expect from such an experiment are shown in Figure 34a. The dissolution rates would exhibit an exact linear dependence on the quantity 6s. For the second experiment, the materials used in the first idealized experiment are subsequently exposed to undersaturated seawater for a period of time. Following the exposure, the instantaneous procedure of the first experiment is employed again. Exposure to undersaturated seawater should increase the solubility of the aragonite surface, leading to the elevated dissolution rate curve shown in Figure 34b. Note that the elevated curve implies that dissolution would be possible in seawater calculated to be

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Figure 34 -a. Idealized experiment in which aragonite dissolution rates are determined instantaneously over a continuous undersaturation range, showing a theoretical linear relationship between dissolution rate and 6s. b. Idealized experiment in which the dissolution kinetics of aragonitic material previously exposed to undersaturated seawater are determined instantaneously. c Idealized experiment illustrating the effects of rate determination at discrete levels of undersaturation, causing enhanced dissolution rates due to the increased surface area of the partially-dissolved surface.

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w a: et: a Experiment. #1 50 DELTA S c 50 DELTA S 1 2 8 b 0 +_ 5(1 DELTA S

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129 when the value of K'spa is used. Thus, the curve in Figure 34b shows a measurable dissolution at 6s = o. This concept is by the observation of increased dissolution rates in the near-saturation range, for perfonmed near the end of a long-term sequence. In the third experiment, which is analogous to my actual experimental procedure, dissolution rates are determined at discrete undersaturation values, and reaction for a period of time. This allows the development of surface dissolution features which increase the active surface area of the aragonite shell. The increased surface area enhances the observed dissolution rate at higher degrees of undersaturation. Because the dissolution rate will be enhanced at each successive level of undersaturation, a curved relationship similar to Figure 34c will be The slow increase in the solubility of the shell surface will also enhance dissolution rates over the length of the Thus, surface can induce an apparent exponential between aragonite dissolution and 6s. However, in the absence of the surface chemical influence, I would expect to observe a linear relationship between dissolution rate and 6s. It is important to note that surface chemical changes will progressively as a shell dissolves in the ocean. As a shell sinks, dissolution will commence in seawater which is only slightly undersaturated. As the shell sinks, it will be exposed to increasingly higher degrees of The effects of surface previously described will accompany the exposure to increasingly undersaturated conditions. Therefore, the process of dissolution

PAGE 156

130 which occurs as a shell sinks is analogous to a slow, continuous multiple-pressure dissolution experiment These experiments have demonstrated an apparent exponential dependence of the dissolution rates on the quantity 65. Since the changes which induce the observed exponential dependence of dissolution rates on in the experiments should occur similarly in nature, an exponential relationship should be used to model the natural dissolution process

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CHAPTER 6: EVALUATION OF RESULTS "Science moves, but slowly slowly Creeping on from point to point." Alfred, Lord Tennyson, Locksley Hall Comparison to Previous Work Several investigations of caco3 dissolution kinetics have preceded the research described in this dissertation. These 131 investigations have used several different techniques and have determined Caco3 dissolution kinetics for a variety of conditions and environments. Although a comprehensive review of all the research in the field is not possible here, it is important to examine how our results relate to the results of previous experimental determinations. Prior to this examination, several reviews of Caco3 dissolution kinetics should be cited. Two important comprehensive reviews are Plummer et al. ( 1979) and Morse ( 1 983) Also, Morse and Berner ( 1979) provide a useful discussion of calcite and aragonite dissolution kinetics in seawater, and the effect of the kinetics on marine sediments. Morse et al. ( 1979) which describes the dissolution kinetic s of aragonite, also summarizes and compares the results of previous laboratory and in-situ determinations for aragonite and pteropod shells. In the INTRODUCTION, initial research on calcite dissolution was

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132 described. Most of this research used the rate equation based on (1-n), Equation 17, to describe the observed dissolution behavior. Although some researchers have used other equations to describe calcite dissolution kinetics, the majority of investigations with aragonite have conformed to the form of Equation 17. The results of the research on aragonite will be compared to our results. Morse et al. ( 1979) examined the dissolution kinetics of synthetic aragonite, and also performed one experiment with pteropod shells. They used the pH-stat method to examine the dissolution kinetics, and found that Eqn. 18 described the kinetics for values of n between 1 and 0.44. The value of n in this range is 2.93. For n values less than 0.44, the value of n increased to 7.27. However, the experiment with pteropod shells indicated that the shells dissolved at approximately "3 percent of the rate of the synthetic aragonite when the dissolution rates are nonnalized to surface area." (Figure 6) The researchers state that the reduced rates are due to the fact that only a small fraction of the surfaces are directly exposed to seawater. Morse et al. also noted that as dissolution proceeded, the ratio of the measured rate of dissolution to the initial rate of dissolution increased. Morse used a pteropod assemblage, and his explanation for the observed increase was increased exposure of surface area which a c companied shell fragmentation. As shown in Figure 35, at approximately 60% dissolution, the rate has increased by a factor of 1.2 over the initial rate. (The curve for calcite is believed to decrease due to preferential dissolution of the small size fraction first.)

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1.2 RM 0.8 RI 133 Calcitic Pacific Ocean Sediment 20 40 60 100 0/o Dissolved Figure 35 (from Morse et. al., 1979) The ratio of measured rate of dissolution per gram to the initial dissolution rate as a function of extent of dissolution for pteropods and calcitic Pacific Ocean sediment (Morse 1978).

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M:>rse et al. 's observations may also be caused, at least partially, by the processes which I have delineated in this study. As shell dissolution proceeds to the point at which 50% of the 134 starting materials have been dissolved, the surface chemistry of the remaining 50% will be significantly altered from the original materials. The degraded surfaces of the substantially-dissolved pteropod shells should lead to enhanced dissolution rates. Also, although lesser exposed surface area on the pteropod shells (compared to synthetic materials) is one important factor, Morse et al. 's results indicate that the dissolution rate of pteropod shells (0.05 is initially much less than synthetic aragonite 2 (1.6 /hr, both measurements at 1-n 0.8). These rates show that small aragonitic particles dissolve at a much greater rate than larger forms, which indicates how the formation of roughened surface features on pteropod shells can lead to accelerated rates of dissolution. Morse et al. (1979) also compared the results of their investigation to the results of in-situ determinations. Berger (1967) and Honjo and Erez (1978) suspended pteropods in the water column, and Honjo and Erez also used synthetic aragonite. Milliman (1977) used aragonitic ooids in a similar manner. Fairly good agreement was obtained between Morse et al. 's results and the experiments of Berger (1967), using H 2o2-treated However, the results of Milliman and Honjo and Erez did not agree well with Morse et al. The main reason for the difference is believed to be the materials used in the experiments. Honjo and Erez used larger pteropods, and Milliman's

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135 ooids should also have been much larger than the pteropods in the >125 ll!ll fraction. The shells used in our experiments were also those from larger pteropod species, ranging from 8 to 45 milligrams. The lower rates of dissolution in this study, ranging from 1 to 5 percent per day, are believed to result from the large size of the shells, and the correspondingly lower surface area/mass ratio of the large shells. Keir (1980) pH-stat dissolution experiments with biogenic calcite and aragonite, using a stirred-flow reactor. Keir's experiments found that the value of the exponent in the rate equation (Equation 20) for pteropods was 4.2. Keir also used synthetic aragonite in his investigation, and found that the rate constant (K) was much less for the biogenic material, compared to the synthetic material. For the synthetic aragonite, the value of K was 13,200; for pteropods, it was 318, even with the presence of broken "chips" in the pteropod experiments Keir' s experiments with various biogenic calcites also showed an increasing rate constant with decreasing particle size. Walter and Morse ( 1985) investigated the dissolution kinetics of shallow marine calcites and aragonites. All of their biogenic aragonites had an exponential value (rate order) approximately equal to 2.5. The samples were crushed to produce grain sizes between 37 and 125 llill For two samples of aragonite obtained from the gastropods Strombus and Hippopus, the exponents were 2.54 and 2.50, and the . value of the rate constants were 302 and 112, respectively, when the results are expressed in terms of percent per day. Walter and Morse did not observe a "discontinuity in the kinetic behavior" as was

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136 observed by Morse et al. (1979) for h i ghly undersaturated conditions. They believe that the discontinuity was related to the use of synthetic aragonite. Walter and Morse also recalculated Keir's results based on new solubility data for aragonite. This recalculation reduced the value of the exponent found by Keir to 3.5, which was still greater than the values found by Walter and Morse. When the results of our experiments were examined as a function of the saturation index (1-n), the value of the exponent n exhibited a wide range of values for different sets of experiments. However, the average value o f the exponent ( 1.789 ) is lower than the value found in previous investigations. Our results now appear to indicate that ( 1-n) should be used only when dissolution is measured under constant pressure conditions. We have not, however, determined whether (1-n) 20[" is appropriate when [co3 Js is held constant. However, if the pH-stat results of several researchers (Morse; Keir; Walter and M orse) are interpreted in light of the results of this study, it now appears that the two methods may illuminate different aspects of cac o3 dissolution The pH-stat m ethod allows samples t o dissolve until a "steady-state" dissolution rate is a c hieved. This method will incorporate processes which enhance dissolution rate, such as the increasing solubility of partially-dissolved surfaces, until t h e apparent steady-state is reac hed. As we have observed that the rate order of the dissolution rate equation increases with enhanced dissolution rate (primarily a result of the data analysis method), the pH-stat method would be expected to lead to greater apparent rate order-s (i.e. increased values of the exponent n). Furthermore, the majority of the pH-stat

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137 investigations have used synthetic which significantly than the shells used in study. The sizes could also to enhanced dissolution due to surface surface areas of the materials will be evident in the value of the constant k. As the dissolution of Caco3 sediments takes place constant conditions in the ocean, and is a process which generally occurs over a long of time (months to the pH-stat method more applicable to the dissolution of caco3 sediments. Our method is more to the initial stages of dissolution of single pteropod shells as they sink in the ocean. In fact, the increasing used in the mimics the that will be encountered by a sinking shell. Our has been confined to a relatively small of conditions that define the dissolution kinetics of in from dissolution rates lead to changes in the mechanism. The investigations of Plummer and associates, and Chou and Wollast et al., 1978, 1979; Chou, 1987), have investigated calcite dissolution a much of solution conditions. The of to the of these investigations is despite the use of in Both also investigated the influence of the dioxide (Pco 2 ) on the dissolution of calcite. The was conducted with pH-stats and one

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138 experiment (performed by Plummer's group). All of the investigations were conducted with essentially pure-water systems, in contrast to our use of seawater. Chou and Wollast used a new method incorporating a fluidized bed reactor in their research. Plummer et al. (1978, 1979) does not provide results which can be directly compared to ours, as the ranges of dissolution rates and saturation state do not coincide. However, Chou (1987) was able to extend the range of saturation states, allowing a comparison to Plummer's data and the results in this dissertation. Specifically, Chou provides values of the molal concentrations (m) of calcium and carbonate ions (Ca2+, CO 2-) in the region near saturation with 3 respect to calcite. Chou's calcite dissolution rate curve and Plummer's curves are shown in Figure 36. The small square shows the data region near saturation where (1-n) is in the range of 0 to 0.5. This data is shown plotted against (1-n) in Figure 37. Chou models the entire data set with the following equation: (43) Rate -8 2 where k 1 = 0.090 em/sec, k 3 = 6.4 x 10 mmol/cm / sec, and k 6 = 15.3 cm4/mmol/sec. In the near -saturation region (above pH 8), the first term in the equation is negligible. Since the third term expresses the effect of calcite precipitation, at equilibrium the following expression is applicable: (44) If the first term in Chou's equation is assumed negligible, the

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139 RAT E OF DISSOLUTION OF Co CARBONAT E 4 ... I -5 0 I N I E -6 0 0 E E -7 ., 0 I) 8 0 -1 9 2 6. Calcite A 6 pH i:x 0 .96 abn C02 0 3 abn C02 0 0 atm C02 X :1< ;l{ L5J '4>. 8 10 X Calcit e 8 Figure 36 -Lei Chou's plot of dissolution rate vs. pH, shown with curves from Plummer et. al. (1978). The square shows Chou's data in the near-saturation region which are comparable to the results in this dissertation.

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,.. (l.l .,.) I1J a::: 0 .... ::) ..... 1 1ft .... 0 2 .4 6 140 Rate units: ( mmo 1 /cm2 /sec ) X t0S 1-0MEGA
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following conversion may be performed (Equation 45): -1 k6 Rate -1 k6 (k3 -k6mca2+ mco32-) = k3/k6 -(m 2+ m 2 ) ca co3 -which is equivalent to ( 46) Rate' 141 If a constant value of mca2+ is then assumed, a condition appropriate to seawater, the expression then becomes (47) This expression is very similar to the equation used to model our dissolution rates in seawater. However, Chou's model is linear, and our model contained an exponential term. Figure 37 shows that the data could be fitted to either a linear or exponential model. Chou's work, while clearly expressing the kinetics of calcite dissolution over a wide range of conditions, is not precise enough to quantitatively define (e.g. linear or quadratic) the nature of the dissolution rate dependence in the near-saturation region. This is shown by the the wide range of dissolution rates at similar levels of undersaturation, as shown in the figure. However, Chou's model agrees with the first conclusion of the previous chapter, that aragonite dissolution rates will exhibit a linear dependence on the quantity (in the absence of other factors which will affect the dissolution rate). Thus, my investigation of dissolution rates, in the near-equilibrium range of conditions which are characteristic of seawater, is in accord with Chou's kinetic

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142 expression. In the near-equilibrium range, the slow rate of the forward (dissolution) reaction is nearly balanced by the rate of the back (precipitation) reaction. The hypothetical rate-determining step under these conditions is the detachment of the ions from the crystalline lattice, in agreement with the mechanism proposed by Morse and Berner (1974). One other aspect of the dissolution kinetics of pteropod shells is worthy of mention. In the models of shell dissolution during sinking, a larger rate constant K was used, to simulate the materials which will be most affected by dissolution, small pteropod shells. effects of dissolution on small pteropod shells, which would be most affected by dissolution. Keir (1980) observed that the rate constant K was inversely correlated with grain size. Keir's rate constant for synthetic aragonite is forty times larger than the rate constant for pteropod shells. Our rate constant for small shells is approximately 3 times greater than the constant determined for large Cuvierina columnella shells. Keir's results support our use of higher values of the rate constant in the model of oceanic dissolution for small pteropod shells. Data Quality The initial analysis of the data from our experiments examined the dissolution rates as a function of (1-0) This analysis provided a data distribution that exhibited a large amount of scatter. Our subsequent use of eliminated much of the scatter, but examination of the datasets indicates that there is still some variability in the

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data quality. The possible reasons for this variability will be examined here. 143 Virtually all of the previous research was perfomed on assemblages of calcitic or aragonitic materials. My research was unique in that it examined the dissolution kinetics of single pteropod shells. One reason for this procedure was to have a constant surface area in each dataset. Because of the different shells, combination of the datasets for a "comprehensive" examination cannot be performed. Different shell masses change the relative amount of dissolution each shell undergoes, even though several experimental sequences were perfomed with shells that had roughly similar surface areas. The primary reason for scatter in the data was the changing activity of the shell surfaces. For example, in our 9-day, 10 experiment sequence, it was difficult to analyze the entire dataset due t o the constantly increasing dissolution rates with time. In other cases, such as the with the data shown in Figures 17, 18, and 31, data points exhibiting substantially elevated rates were eliminated from the dataset to allow least-squares analysis of the curve. Furthermore, as the importance of the surface alteration effect became evident, the preparation and storage procedures were adjusted so that surface alteration would be limited as much as possible. Figure 28 showed the effect of surface alteration when a highly-undersaturated storage solution was Certain aspects of at-sea performance of the research made the data quality obtained at-sea remarkably good. The main reason for the high data quality was essentially continuous e x perimentation. Shells

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were not stored between experiments, but were used in consecutive experiments until retirement. 144 Another aspect of the research at-sea was the availability of different pteropod shells for the experiments. Shell size and morphology made some species easier to use in experiments than others. Cuvierina columnella is ideal for the experiments, as its shape allows it to move freely in 'the sample chamber. Shells such as Cavolinia tridentata and Clio pyramidata occasionally became stuck in the chamber, and their overall movement was somewhat restricted. Instead of falling freely through the chamber during each oscillation, the shells of those species could slide down the sides, which would inhibit exchange of seawater under the shell, and limit stirring of the contents. The primary mixing mechanism in our experiments was the movement of the shell within the chamber. These limitations may seem trivial, but the high sensitivity of the technique could accentuate these variations. The "kink" and "step" theory of d issolution (Burton and Cabrera (1949), Burton, Cabrera and Frank (1951); cited by Morse and Berner (1974) and Morse (1983)) has ion detachment occurring at specific sites on a crystal surface. These sites are commonly found on edges and flaws in the shell or test. For this reason, homogenous crystal assemblages would be expected to dissolve in a more regular fashion than irregularly shaped shells. After surface area and solubility changes have been limited by proper experimental procedures, the primary cause of the remaining "noise" in the datasets is the surface variability of the individual shells. When all of the possible sources of experimental noise are

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145 limited, the system is capable of producing high quality data. An example is Figure 26, the Cuvierina columnella data used to analyze the value of 6. V for aragonite, and dataset 14 (Table 3) an experimental sequence performed with the best preparation and storage procedures. These datasets should not be viewed as exceptions, but as illustrations of the system's capability under optimum conditions.

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CHAPTER 7: ARAGONITE SOLUBILITY AND THE OCEANIC SYSTEM "Though they go mad they shall be sane, Though they sink through the sea they shall rise again; Though lovers be lost love shall not; And death shall have no dominion Dylan Thomas, And death shall have no dominion. ----..;......;.._;;_.;,_....;.._ Investigation of in-situ Dissolution with Laboratory System One of the advantages of the high-pressure system is the 146 simulation of actual oceanic conditions of temperature and pressure without requiring a visit to oceanic depths. The system has been used to confirm chemical determinations of the saturation horizon in the North Pacific Ocean. Also, a profile of dissolution rates with depth indicated that the dissolution rate of aragonite increased with depth, despite lesser values of the quantity (1-G). The spectrophotometric determination of pH has also been used at sea to measure the pH of seawater samples, which were compared to standard potentiometric measurements (Byrne et al., in press). In May of 1982, the potentiometric system was used to confirm the depth of the aragonite saturation horizon at 49.5'N, 175.0'E in the North Pacific Ocean. Aragonite fragments and the shells of three different pteropod species were used at 3.6C and pressures approximately equal to in-situ pressures at 123 and 418 meters. The

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147 in-situ temperature at 123 meters was 3.64C, and at 418 meters was 3.42C. Then value for aragonite at 123 meters was about 0.7, and at 418 meters about 0.55. Our potentiometric measurements indicated that dissolution occurred at both depths (Feely et al., 1984). In June of 1985, the spectrophotometric system extended the scope of this application at 50N and 145W in the North Pacific Ocean. The system was used to measure the dissolution rate of a single Cavolinia tridentata shell at 5C and pressures corresponding to depths from 100 to 3000 meters. The location was Station 13 of the research cruise aboard the R / V Discoverer from Honolulu, Hawaii to Kodiak, Alaska. As mentioned previously, this data set was also analyzed as a function of for aragonite dissolution. One of the interesting characteristics of the saturation state of the water column in this region is the Q minima which occurs at either 1500 or 2000 meters. Q increases below that depth to 3000 meters. This profile was exaggerated in our experiments, as all the experiments were performed at 5C. The higher temperature caused the n value to be even greater than in-situ conditions in the highest pressure experiment. (Larger values of n, up to a value of 1, indicate lesser degrees of undersaturation.) Figure 38 shows the dissolution rates that were obtained at 100, 125, 150, 175, 200, 250, 350, 500, 1000, 1500, 2000, and 3000 meters. Although Q reaches a minimum value at 1500 meters and then increases, the dissolution rate increases with depth. This observation was one of the first indications that dissolution rates were not exactly correlated with (1-Q) under conditions of variable pressure. Figure 38 represents the first detailed profile of actual aragonite

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38 -Dissolution and Q at S0N, 14S0W in the Pacific Ocean. Each dissolution was obtained using a single f tridentata shell, mass 16.2 1 mg. The calculated Q profile corresponds to the saturation state of our shipboard experiments which performed at SC. corresponded to the depths which the samples obtained.

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I 1-a.. w 0 RATE (%/DAY) 1500 DISSOLUTION RATE 2700 vs. DEPTH 50N 145W Ca volin/a trldentata 16.129mg 3000L--------------------------ARAGONITE SATURATION PROFILE 1 0 0.9 0.8 0 7 0.6 0 5 0 ARAGONITE 149

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1SO dissolution rates under representative oceanic conditions. Earlier in the cruise, water were obtained from depths that were close to the aragonite saturation horizon in the region. Samples were obtained at stations 8,9,10, and 11. The in-situ water temperature of these samples ranged from 6. 9 to 4. 1 c, and was generally within 1.0C of SC. These samples were used to measure the dissolution rates in the vicinity of the saturation horizon depth at each of these stations. From station 8 to station 11, the saturation horizon rose approximately 100 meters. Figure 39 shows the dissolution rates measured at these stations, and the 100% and 80% (Q = 1.0 and 0.8) saturation horizons. The figure shows clearly that dissolution increases rapidly below the saturation horizon in this region. The small rate of dissolution measured above the saturation horizon at station 9 may be the result of a combination of factors. At this depths, the in-situ temperature was higher than SC. Various shells were used in these experiments, as they were performed over a period of time during the research cruise. This figure and the dissolution rate profile (Fig. 36) demonstrate that the system is capable of useful assessments of the natural saturation state of the ocean. The capability of the system to observe dissolution under very slightly undersaturated conditions is one of its strongest features. Because of the differences in the value of K 'spa from various experimental determinations, uncertainty in the value of K'spa could make a substantial difference in the calculated depth of the saturation horizon. Direct confirmation of the depth of the saturation horizon indicates that the value of K'spa determined by

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-E ......... I t w 0 Dissolution Rates in %/day Station 8 9 10 II 47N O .OJ .01 300 n Aragonite(%) 400 500 600 151 Figure 39 Plot of measured dissolution rates at four stations in the North Pacific Ocean, July 1985. Experiments were performed at near in-situ pressures and 5C. The in-situ temperature range was 6 .75 to 4 1C. Measured rates coincide with the calculated saturation horizons, increasing rapidly below the 100% saturation horizon

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152 Morse et al. (1979) and Mucci (1983) is accurate. The system also demonstrates that dissolution of aragonitic shells should begin quite rapidly after the shells descend below the saturation horizon. Previous models (Berner and Morse, 1974; Broecker and Takahashi, 1977) predicted that significant dissolution would begin only after a critical level of undersaturation was reached. This recent data indicates that dissolution will commence immediately upon exposure to undersaturated seawater. However, the initial rate of dissolution in seawater which is only slightly undersaturated is very slow. Use of the Saturation Parameter in Oceanographic Research Many investigations concerning Caco3 dissolution kinetics in the ocean have been directed toward quantifying the magnitude and extent of the dissolution process in the ocean. Oceanic absorption of atmospheric co 2 caused primarily by increasing co 2 concentrations in the atmosphere from fossil fuel combustion, will decrease the depth of the saturation horizons and increase the degree of undersaturation in the water column. However, Caco3 in the oceans may act as a buffer for the increased amount of co 2 that is being added to the atmosphere, and which is subsequently absorbed by the ocean. As aragonite is the most reactive of the two dominant forms of Caco3 found in the marine environment, the processes affecting its fate will be involved with the present-day co 2 (Calcite with a high magnesium (Mg) content, greater than 12 mole percent, is more soluble than aragonite

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153 in seawater. High-Mg calcite is formed in shallow-water regions. Transport of this material off reefs and banks will allow some dissolution in the deep ocean, but the contribution of this material is believed to be substantially less than either aragonite or calcite.) Previous investigations of Caco3 dissolution kinetics have correlated the dissolution rates with the relative saturation index (1-n). The results of those investigations have been extended for predictions of Caco3 dissolution rates in the ocean. These results have been successful in determining the approximate depth of the saturation horizons and carbonate (aragonite or calcite) compensation depths. However, this research has demonstrated the applicability of the absolute saturation index ([co32-Js-[co32-J) or to aragonite dissolution rates in the oceanic water column. The experiments performed in this research suggest that is a more useful variable than n for prediction of aragonite dissolution rates under conditions nearly identical to those found in the ocean. For many years, the total alkalinity and total co 2 concentrations in the water column have been measured in order to determine the saturation state of the water column with respect to aragonite and calcite. The calculated saturation state has been correlated with the existence or disappearance of carbonate sediments on the ocean floor. In our initial research (Byrne et al., 1984), we determined that the shallow saturation horizon and high degree of undersaturation in the northern Pacific should lead to substantial dissolution of pteropod shells before they reached the ocean bottom Therefore, the

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saturation state of the ocean will affect Caco3 dissolution even in areas where carbonate sediments do not accumulate on the ocean floor. 154 The relative saturation index Q has been widely used to indicate the oceanic water column saturation state (Edmond and Gieskes, 1970; Morse and Berner, 1978; Feely et al., 1984). In the north Pacific Ocean, it is common to find a minimum value of n in the upper 2000 meters, below which n increases slowly. 63, the absolute saturation index, behaves somewhat differently. M is relatively constant between 2000 and 4000 meters, and has a rapidly increasing trend below 4000 meters. Figure 40 shows two profiles of Q and generated from data collected in the Pacific Ocean in July 1985. The scale has been shifted so that the profiles approximately coincide in the upper water column, where the dissolution rate predictions of both saturation indices are similar. Note the divergence of the two indices as depth increases. Figures 41-48 show cross-sections of 6s and 0 in the eastern and western North Pacific Ocean, the southeastern and equatorial Atlantic Ocean, and the southwest Indian Ocean. The Atlantic data is from the GEOSECS expedition (Takahashi et al., 1980), re-calculated according to Brewer et al. ( 1 983). The Indian Ocean data is from the north-south cruise track of the R / V Marien-Dufresne in February and March of 1985. The Pacific Ocean data is from the cruise tracks of the R/V Discoverer in May-June 1982 and June-July 1985. 6s values in the deep Pacific are the largest values calculated for any oceanic region. is a useful variable for mapping the susceptibility of Caco3 to dissolution in the global marine environment. M makes it relatively easy to assess relative aragonite dissolution rates. It is

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155 Q 8 .6 4 2 b a 1 2 2 \ I I 0
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E .::1! .r= a. 0 156 Figure 41 Plot of 6s in the E. North Pacific Ocean, July 1985. (Unpublished data provided by J. Gendron)

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157 1 eo C0 2 -E -3 .t= a. Q) a Figure 42 -Plot of Q in the E. North Pacific Ocean, July 1985. (Unpublished data provided by J. Gendron)

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..... E ...., .t:3 Q. Q) 0 4 6 158 30 Figure 43 -Plot of 6s in theW North Pacific Ocean, June 1982. (Data from Feely et. al., 1984)

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159 a x100 Degrees Latitude 20 30 40 50 0 300 200 100 100 eo eo eo sO eo 2 -E .._3 eo .t::. -c. Q) 0 Figure 44 Plot of Q in the W North Pacific Ocean, June 1982. (Data from Feely et. al., 1984)

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1 60 Degrees Latitude 20 10 0 -20 0 3 0 Figure 45 Plot of in the S. Atlantic Ocean. (GEOSECS data)

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161 0 x100 Degrees latitude 60S 50 40 20 10 0 0 1 e;.. Qoo 100 100 80 Figure 46 Plot of Q in the S Atlantic Ocean. (GEOSECS data)

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...... E X ._, 162 6s 50 Degrees Latitude 40 30 -100 -eo -----------------1 1 -20 ----------------1 0 4 Figure 47 Plot of 6s in the S. Indian Ocean, February-March 1985. (INDIGO I data)

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2 ...... E :.e. ..... Q. a> 0 4 50 0 x100 40 100 100 80 163 Deg rees Latitude 30 8 0 Figure 48 Plot of n in the S. Indian Ocean, F ebruary-March 1 985. (INDIGO I data)

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164 clear from the 8s cross-sections that below 4000 meters the waters of the North Pacific should be highly corrosive to aragonite. A desirable extension of this work would be to correlate the values of with geological markers of dissolution in aragonite-dominated sediments. Further research is required to determine the relationship between calcite dissolution rates and c (for calcite). Because calcitic forms are generally more irregular and porous than aragonite shells, surface chemical changes will be less likely to measurably influence the dissolution kinetics of these forms. The effects of breaching and partial disintegration are likely. to be the dominant processes which will alter the dissolution rate of biogenic calcite materials. The majority of carbonate sediments in the ocean are composed of calcite, and geological markers of dissolution for foraminiferal assemblages are already well known. Since these geological markers are already correlated with depth, and the carbonate chemistry of the water column has been measured in major oceanic regions, it would be relatively simple to calculate and correlate the parameter with sedimentary dissolution indices. Processes Affecting the Dissolution of Aragonite in the Ocean ---Laboratory investigations of aragonite (or calcite) dissolution tend to analyze the process in ways that limit the variation of experimental conditions. By maintaining uniform solution conditions and varying only a single parameter of interest (such as n or phosphate concentration), these techniques overlook the highly dynamic

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165 nature of dissolution as it takes place in the The results of initial on aragonite dissolution in the ocean et al., 1984) noted how dissolution would affect the settling of pteropod shells in the water column. In the reduction of settling caused by mass loss due to dissolution led to a prediction of exposure to undersaturated deep waters, which in the rapid of settling shells. More recent has indicated several processes which will accelerate dissolution as a shell sinks to the ocean The following discussion delineates these processes The determining the dissolution of debris is the dependence of the dissolution rate on the saturation state of the water mass (which is determined by the pressure, and salinity). As stated several investigations have described the dissolution of aragonite and calcite as functions of the relative saturation index (1-n). The work described in this indicates that the absolute saturation index ([co32-Js [co/-J) is a more index of dissolution kinetics. 40 showed that b,s increases with depth in the colwnn, but (1-Q) constant. Therefore, b,s indicates that dissolution will continue to with depth, which is not apparent the values of (1-n) with depth. Data from 35N, 165E in the North Pacific Ocean (Feely et al., 1984) show that b,s by a factor of 26 ( 3. 2 to 85.27) the saturation horizon to 6000 meters depth. In contrast, (1-n) by only a factor of 12 (0.044 to 0.547) at the same location. This

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contrasting behavior indicates that dissolution rates predicted by models based on should be greater or equal to dissolution rates predicted by models based on ( 1-n), especially for deep oceanic conditions. 166 The actual dissolution rates measured in our experiments are somewhat smaller than the rates measured in previous investigations. The maximum rates of dissolution found in our experiments were no greater than 5% per day. Previous investigations, including Byrne et al. 1984, found maximum rates ranging from 30-50% per day. The primary reason for this difference is the large pteropod shells used in the spectrophotometric system. Large shells provided fairly large dissolution signals, but due to their size, the percentage loss to dissolution was small. However, small pteropod species dominate the size distribution of pteropods in the ocean. Species such as Limacina inflata and Limacina retroversa are found in large numbers in some areas of the ocean. In these regions, the pteropods can even be a food source for fish and whales (Be' and Gilmer, 1977). Be' and Gilmer use the terms "ubiquitous" and "abundant" to qualitatively describe the numbers of pteropods in these regions. Limacina inflata has been collected in densities exceeding 100 organisms per cubic meter (McGowan, 1960), and one of two Limacina helicina subspecies had a maximum density of 76 organisms per m3 (McGowan, 1963). The shells of these small species generally weigh much less than milligram. Typical shells of L. retroversa weigh less than 0.1 milligram. In the rate equations (equations 17 and 36), the value of the rate constant K is dependent on the saturation index used in the

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167 equation, and on the type of used in the dissolution investigations. Because the smaller shells and K these materials should be leading to ve of dissolution the smaller shells. Cumulative Affecting Dissolution with Sinking It is now to examine the impact that the discussed above will have on the dynamics of dissolution in the ocean. As a pteropod shell sinks, it will go from a relatively high-temperature, into an of increasing and decreasing It is well known that the solubility of aragonite these conditions The quantities and n indicate increasing with depth in the water column, with showing a with depth very high All of these will cause the dissolution of a shell to accelerate as it sinks into deeper as the shell sinks and begins to dissolve, it mass loss which slows its of descent. on the settling rate is the increased viscosity of seawater at low temperatures. in indicated that settling velocities of pteropod shells by fifteen 0 0 as the 25 C to 5 C. As the shell's rate of descent the shell's time to rapidly Other that as the shell dissolution

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168 the increase in the surface area of the partially-dissolved shell, and a possible slight increase in the solubility of the shell surface. Thus, as the shell sinks, it will be subjected to several processes which may combine to increase both the dissolution rate and the exposure time to undersaturated conditions. As all of these affect the shell during its descent, it is clear that dissolution can dramatically affect the shell as it sinks. Modes of in which the shell is exposure to undersaturated seawater, similar to fecal pellet of coccoliths, have not been described for pteropod shells. Refined Model of Shell Dissolution with Sinking In et al. (1984), a model of shell dissolution with sinking was developed. This model indicated that slowly-settling shells would before the ocean in the northern Pacific Ocean. One of the of the model was the of the sinking rate that as the shell's mass However, no other sinking rate which accelerated the dissolution were incorporated into the model. Furthermore, the model divided the ocean into ten 500-meter depth layers, which did not sufficiently indicate all the features of the changing water column state. The results generated by this model have previously been shown in Figures 15 and 16. A new model of shell dissolution which all of the processes previously including the reduction in shell mass which accompanies dissolution, has been developed. This model is

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169 based on the saturation parameter, and it divides the ocean into any number of layers based on the available data. Also, once the temperature decreases below 5C, the settling velocity is reduced by an additional fifteen percent due to the increased viscosity of the seawater. (Measurements of shell sinking velocity with decreasing seawater temperature were performed in the laboratory. Table 6 shows the data which was used to estimate shell settling rates.) This model is highly sensitive to the value of K and n in the rate equation. In order to investigate the effect of varying these parameters, the model was used to generate several predictions of cumulative dissolution for various settling rates, values of K, and values of n. The value of n from our "best" set of data (.f.. columnella, Figure 26) was 1.81, for a value of -37.0 cm3 /mol. Based on our previous discussions, the value of K for small shells is believed to be greater than the value of K for the large shells used in our experiments. For our initial investigation of the dissolution model, a value of .01 was chosen, so the form of the rate equation is ( 48) RATE (%/day) 0 0 1 ( ) 1 81 for the first predictions of the model. Figure 49 shows the effect of variations in the value of K and n on the predictions of the model. Curve I shows the results of the rate equation shown above, for a shell settling with an initial (25C) settling rate of 1 .5 em/sec. The model predicts that the shell would be completely dissolved at about 5500 meters.

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TABLE 6 Data on pteropod shell settling rates. Each experiment will be described separately, as several different methods were employed Experiment 1 -A single Diacria trispinosa was used in a 75c m experiment conducted a walk-in freezer. The reduction in speed with lower temperature is demonstrated. Seawater from the Gulf of Mexico with a salinity of 36.677 00 was used. Temperature (oC) Time (seconds) Rate (em/s ec) 17.8 25.5 2 2 .94 16. 8 24.60 3.05 12.0 25.94 2.89 9.9 25.64 2.93 9.0 26.17 2.87 7.0 26.35 2.85 4 8 28. 1 3 2.67 2.8 28.89 2.60 1 4 29.85 2. 51 170

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171 TABLE 6 (continued) Experiment 2 -Two small L imacin a i nflata shells were used in a 20-cm test tube filled with 36.677 /00 salinity seawater. The timings were performed at 20C, and then the t ube was cooled in ice to 2c Pteropods were released from a glass pipette before timing began. (Times are in seconds settling rates in centimeters/second.) "Fast" L. inflata 20C -Time Rate 2C -Time Rate 7 .75 1.29 10.05 0.99 7 .85 1.27 8 .01 1.25 7.36 1.36 9 5 1 1.09 7.56 1.32 8.25 1.21 7 .18 1.39 8.10 1 .23 "Slow" L. inflata 20C -Time Rate 2C -Time Rate 8 .74 1.14 8 41 1.19 8 .25 1.21 8.80 1.14 7.99 1.25 10.74 0.93 8.39 1.19 11.34 0.88 7 .22 1.38 10.28 0.97 8 .72 1 .15

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TABL E 6 (continued) Exper iment 3 Several different species of pteropods wer e t imed over a 1 meter transit at 23C, i n seawater o f 35.28 00 salinity. L i macina bulimoides : Time (seconds) Rate (em/sec) 4 specimens Tr ial 1 : 33. 9 4 2 .95 40.15 2 .49 44.10 2 .27 40. 9 1 2 .44 Trial 2: 36.99 2 .70 36.52 2 .74 37.03 2 .70 36.33 2 .75 Cuvierina columnella : 17.75 5.63 3 specimens 16.65 6 .00 16.36 6 1 1 Clio cuspidata: 41 .27 2 .42 Limacina helicina: Specimen 1 53.10 1.88 Specimen 2 62.69 1.60 S p ec imen 3 (tw o runs) 83.02 1.20 82.72 1. 21 172

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173 TABLE 6 (continued) Experiment 4 Several individual Limacina retroversa and Clio pyramidata were used in a 31 4 em transit measurement. A graduated cylinder was placed in the walk-in freezer, so that measurements took place below 5C. Each measurement represents a different shell, and the results are shown chrono l ogically. Species T t (oC) empera ure L. retroversa 4.7 4.0 3.5 3.2 pyramidata 2 4 1.7 1 2 L retroversa 0.9 0.5 Time (sec) 19.70 19.89 18:14 16.91 15.44 25.56 22.3 1 35.31 19.75 17.20 8.74 10. 11 7.49 10.83 15.44 9 .28 7. 61 6 .47 12:12 7.55 15.29 14.04 20.4 2 15.55 16:88 30.18 Rate (em/sec) 1.59 1.58 1. 73 1:86 2:03 1:23 1 41 0:89 1.59 1:83 3 .59 3.11 4:19 2 .90 2:03 3 .38 4:13 4 .85 2:59 4.16 2 .05 2 .24 1:54 2 02 1.86 1.04

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1 :c .t3 0.. w 0 4 5 k .01 .004 n IV: -37.0 cm3/MOLE 1 00 80 60 40 20 0 %OF SHELL REMAINING 174 Figure 49 Model of dissolution with sinking, using data from 50N, 167W in the North Pacific Ocean. Dissolution rate equations used to generate curves: I. Rate = .01(6s) 1 81 II. Rate= .004 (6s) 1 .81 III. Rate = 0 1 (6s) 1 5 IV. Rate = .01 (6s) 1 81, 6V for aragonite dissolution= -37.0 cm3 1rnol. 6V = -31.3 cm3 1mol for the first three curves.

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175 If the value of K is by a of 2.5 II is in dissolution in of the shell in the column. In this case, at 50N in the western Pacific Ocean, 50% of the shell would the sea floor. The predictions of the model are even more sensitive to the value of the exponent n. The value of the exponent was 1 .81 to 1.5, and K was to the value, .01. In this case, the increasing of has little effect on the shell, and almost 80% is to the sea floor. Curve IV is with the value of 6.V equal to -37.0 cm3 /mol, than -31.3 cm3 /mol, which was used for the other Because a higher value of 6.V a of elevated the depth to which the shells descend complete is shallower. Based on dissolution noted on shells caught in sediment traps deployed in the Pacific it seemed useful to attempt to the of the model. In 1982, we found several Limacina inflata shells at 35N, 165E in a deployed at 2170 meters. These shells substantially affected by dissolution at this fairly shallow depth. SEM examination of the shells et al., 1984) showed the microcrystalline which the of dissolution. Small Limacina inflata shells are especially in this discussion as they settle at 1 .0 1. 5 em/sec. The model based on rate equation 48 above that shells settling at 1 .o em/sec would lose 20% of mass by 2538 meters, to shells

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176 settling at 1 .25 em/sec, which would lose more than 10% of their mass in transit to this depth. The prediction of the model therefore accords with the indications of surface dissolution observed on the shells. Adelseck and Berger (1975) collected pteropods and other settling Caco 3 detritus by Bongo net tow at 3000 meters in the eastern tropical Pacific. An SEM microphotograph of a Limacina shell (-500 diameter) shows almost complete surface etching over the entire shell. Data is not available for this area, but the aragonite saturation horizon is probably deeper than in the northern Pacific. Furthermore, a shell caught in the net would have been exposed to undersaturated conditions only during settling, and not for an extended period of time in a sample chamber. Thus, the shell shown in Adelseck and Berger (1975) would have been exposed to undersaturated seawater for a minimum period of time. Another location that allows rough correlation of dissolution with the predictions of the model is at "Station 13" of the 1985 Northern Pacific Cruise, at 50N and 150W in the northern Pacific. At this station, both shallow sediment traps and surface plankton tows indicated a large abundance of Clio pyramidata pteropods. At this location, a trap was deployed at a depth of 3500 meters. Several Limacina helicina shells were recovered from this trap. These shells showed substantial indications of dissolution. Clio pyramidata is a large pteropod species (see figure 1), but is relatively thin-walled. Large numbers of this species were collected at the surface in net tows at Station 13, and several specimens were collected in the trap deployed at 2100 meters. These

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177 shells also showed substantial dissolution effects, including almost complete opacification and partial fragmentation. Despite more rapid settling rates than Limacina helicina shells, only a single Clio pyramidata shell fragment was collected in the 3500-meter sediment trap. This observation may imply a more rapid mode of destruction for the Clio shells, but could also be due to species abundances and depth distributions. However, observations of Clio pyramidata shells settling in a cylinder indicate that the shells tend to sink in an orientation that allows a rapid rate of descent, considerably in excess of 1 .5 em/sec. Therefore, for dissolution effects to be apparent on the shells recovered at 2100 meters, the rate of dissolution must be relatively rapid. This statement is valid even if the shells resided in the traps for the maximum possible amount of time possible, 24 hours. A shell settling at 2.0 em/sec would reach 2100 meters in slightly more than one day. Therefore, the maximum amount of time the shell could be exposed to undersaturated seawater is three days. The model predicts that a shell settling at 2 0 em/sec at this location would disappear at approximately 5000 meters in about seven days. It should be apparent that the dissolution rate must be fairly rapid to have a substantial effect in this short period of time and at this depth. Therefore, the predictions of the first rate equation (which generated curve I in Figure 49) appear to agree with observations better than the predictions of the modified rate equations ( curves II and III). It should also be noted that the values o f 6s shown in the transect figures (Figures 41, 43, 45, and 47) have been generated with a value of 6V equal to 31.3 cm3 / mol, so that the results for the North

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178 Pacific data would agree with those of Feely et al. (1984, and submitted ms.). A larger value of 6V would have the effect of increasing the magnitude of 6s with depth, which would indicate even greater dissolution rates at depth, as shown by Curve IV in Figure 49. Figure 50 uses the above rate equation (Eqn. 43) to show the effect of increasing settling rate o n the predictions of the model. This data was for Station 6 in the mid-northern Pacific in 1985. Clearly, the depth at whic h the shells disappear increases with increased settling rate. Note the rapid acceleration in the mass loss that occurs below 4000 meters, as the effects of increased pressure and shell erosion begin to take effect. Note also that the mass loss higher in the water column will slow the rate of descent, so that the shells are exposed to the highly corrosive conditions in the deep water column for extended periods of time. The deceleration also contributes to the rapid mass loss below 4000 meters. Figure 51 us e s a mid-range settling rate (1 .5 em/sec) and water column data from the northern Pacific and Indian Ocean. For the northern Pacific stations, the model indicates that the depth to which the shells will descend befor e being consumed by dissolution lies within a fairly narrow range (500 meters), despite the 500-meter difference in the depth of the saturation horizon. However, the saturation situation in the Indian Ocean is clearly different. The model predicts that only about 10% of the shells will be dissolved when they reach the depth of the ocean floor at this station, and that even in a 6000-meter water column 40% of the shell would survive to r e ach the ocean floor. In this particular region of the ocean, a small, ubiquito us pteropod species, Limacina retroversa,

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1 J: 3 ._ a.. w 0 4 5 100 80 60 40 20 0 %OF SHELL REMAINING 179 Figure 50 -Model results for various settling rates from 1 to 2 em/sec. Equation: Rate = 01(6s) 1 81

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Or---------------------------1 \ \ \ .. STA.13 -STA. 6 ---STA.9 :::C3 t\ . \ \ a.. w \ .. \ PACIFIC OCEAN 04 INDIAN', OCEAN' . 5 .. .. .. .. .. BOTTOM',, ..... ... DEPTH ,__ ............. X ... 50 .____ X ---x 100 80 60 40 20 0 % OF SHELL REMAINING Figure 5 1 -Model results for settling rate of 1 5 em/sec. 180 Equat ion: R ate= Data from Northern Pacifi c station s 6, 9 and 13, (July 1985), and 50S in the Indian Ocean.

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181 is abundant in the surface waters (Be' and Gilmer, 1977; unpublished observations from INDIGO I cruise). The model indicates that there should be a large amount of aragonite in the form of these shells which reaches the ocean floor prior to dissolution. However, at 4000 meters, the value of is approximately 40, which indicates that dissolution on the ocean floor will be relatively rapid. Aragonitic sediments are not found in this region (Figure 2), so the observed conditions are in accord with the sedimentary record. Note that at 3000 meters in the Indian Ocean, very little dissolution of the shells has taken place, and the value of 6s is approximately 10, so the dissolution rate is slow. These factors indicate that sediment traps deployed at this location should give a good indication of the aragonite flux which reaches this depth, and the results would not be substantially influenced by dissolution. This is a region in which the aragonite flux to the deep waters should be large. Therefore, flux data from this region would be a valuable indication of how much aragonite in the form of pteropod shells descends to the deep water colurrm, in regions of high "aragonite productivity". Because of the difficulty of quantifying the aragonite flux in the ocean (see discussion below), and the difficulties of sediment trap sampling in undersaturated waters, research of this nature could help resolve the uncertain magnitude of the aragonite flux. Estimates of the Aragonite Flux in the Ocean As stated previously, considerable uncertainty exists as to the

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182 magnitude of the aragonite flux in the ocean, and its importance relative to the calcite flux. Although many sediment trap investigations have been performed, only a few of these investigations have quantified the calcite and aragonite contributions to the total Caco3 flux. Furthermore, sediment trap investigations may be influenced by biological and chemical factors. Pteropods are mobile, responsive organisms, and their behavior may influence the trapping process. The solubility of aragonite may also cause the selective loss of the mineral when traps are deployed in undersaturated water. In the following discussion, I will attempt to consolidate the available information on aragonite fluxes in the ocean. One reason for this examination is that the original investigations which led to this dissertation were directed at quantifying the aragonite flux. It will be clear that considerable uncertainties remain which make it very difficult to render conclusive estimates. However, it is worthwhile to examine the available information, as it will become apparent that upper-bound estimates of the aragonite flux imply that aragonite dissolution could be a major process in the oceanic co2 system. Models of the oceanic co 2 system have attempted estimating the rates of calcite and aragonite sedimentation. Broecker (1971) created a model of the caco3 budget in the oceans, which was described by Berner (1977). Fiadeiro (1980) modeled the alkalinity distribution of the Pacific Ocean, and compared his results to the known sources of alkalinity in the Pacific. Broecker's original model assumes that the rate of aragonite

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183 sedimentation is negligible, and estimates the average dissolution rate of aragonite plus calcite in an ocean using "independent mass-balance calculations" (Berner, 1977). Broecker's rodel produced a calculated value for the rate of calcite sedimentation in the Atlantic Ocean of 5.0 mg/cm2 / year, and for the Pacific 2.5 mg/cm 2/year. The observed rates of calcite sedimentation are 1 .0 mg/cm2/year for the Atlantic and 0.5 mg/cm2/year for the Pacific. Following Broecker's suggestion, that neglect of aragonite sedimentation might be a principal reason for the disagreement between the observed and calculated sedimentation rates, Berner postulated that the rate of aragonite sedimentation was equal to the rate of calcite sedimentation. When the assumption of equal calcite and aragonite sedimentation rates is used in Broecker's ax>del, the calculated values of the sedimentation rates are much closer to the observed values. The ax>dified "Broecker-Berner" model (Berner, 1977) now provides calcite sedimentation rate estimates of 1.3 mg/cm2/year for the Atlantic and 0.9 mg/cm2/year for the Pacific. Fiadeiro (1980) found that his model of the alkalinity distribution in the Pacific required a source of alkalinity in the mid-depths of the water column, from roughly 2000 to 4000 meters. As calcite particles are assumed to be unaffected by dissolution during settling, and the magnitude of aragonite sedimentation is assumed to be negligible relative to calcite, no apparent source of alkalinity in this region of the water column is available. Fiadeiro' s model 2 2 required approximately 35 mg/m /day, or 1.3 mg/cm /year of Caco3 to provide the necessary Tsunogai and Watanabe (1981) calculate a "carbonate dissolution

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184 flux" in the Pacific Ocean equal to 63 mg/m2 /day, or 2.3 mg/cm2/year. Berner 977) estimates relative abundances of pteropods and foraminifera in the ocean and in sediments, to determine the validity of assuming that the rate of aragonite sedimentation is equal to the rate of calcite sedimentation. Berner cites plankton research performed by Adelseck and Berger ( 1975) and Berger ( 1 976) as indicating that pteropod and foraminifera abundances are approximately equal. Adelseck and Berger used vertical net tows from 0-200 meters to estimate the abundance of foraminifera and pteropods. Three tows found foraminiferal abundances of 8.6, 19.4 and 12.8 individuals per cubic meter. The corresponding pteropod abundances were 11.3, 10.7, and 5.3 individuals per cubic meter. Berger (1978) confirms these observations, stating "In the living plankton the shell mass of pteropods and heteropods is comparable to that of the foraminifera, and can actually exceed it substantially." However, because biological regeneration rates are slower for pteropods than for foraminifera, with pteropod life-spans estimated at 18-24 months and foraminifera 1-2 months, the actual flux of calcite and aragonite from surface waters may to the deep ocean may not be in the same ratio as the surface abundances. Because the turnover rate for foraminifera should be more rapid than the turnover rate for pteropods, a larger number of foraminiferal tests than pteropod shells is expected in the downward caco3 However, it should also be noted that a single large pteropod shell may represent 10 to 100 times as much Caco3 mass as a foraminiferal test. Therefore, despite the fact that foraminifera individuals in the descending flux may substantially outnumber pteropod individuals, the descending mass of

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185 aragonite and calcite could still be roughly equivalent. Sediments found at depths where the water is supersaturated with respect to calcite and aragonite can also provide information on relative amounts of pteropod shells and foraminiferal tests. Berner (1976) determined relative numbers of pteropods and foraminifera in sediment samples obtained on the Bermuda Platform. For the two shallowest samples (the only two above the aragonite saturation horizon) he found the ratio of shells/tests was 1.01 and 1 .05, at depths of 1475 and 1672 meters. As these are specimen-to-specimen ratios and not weight ratios, and pteropod shells are generally more massive than foraminiferal tests, the ratio of aragonite mass to calcite mass is likely to be greater than 1 in these sediments. However, the activities of burrowing organisms have been shown to "buoy" large particles to the top layers of a sediment, which would augment the numbers of pteropod shells compared to foraminifera at the sediment surface (Price et al., 1985). Chen (1964) counted pteropod shells in Bermuda platform cores from the Lamont-Doherty Geological Observatory core locker, and determined that the shallowest cores were over 90% Caco3 and contained over 600 pteropod specimens per milligram dry weight Considered with Berner's pteropod/foraminifera ratios, it is clear that a large percentage of the caco3 in the Bermuda sediments is aragonite. Berner (1977) also cites data provided by Herman, indicating that the pteropod/foraminifera ratio in shallow Red Sea and Mediterranean sediments is approximately 1.0. Therefore, as Berner suggests, it is reasonable to assume that the aragonite/calcite ratio in the Bermuda sediments is approximately 1.0.

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186 The above discussion leads to the conclusion that the mass flux to the deep ocean should be approximately equal to the calcite mass flux. However, as a follow-up to the estimates in Berner (1977), Berner collaborated with Honjo in an investigation of the amounts of and calcite found in sediment and Honjo, 1979). Sediment trap collections analyzed by diffractometry to the proportions of calcite and aragonite in the samples. The determined that aragonite 12% of the total Caco3 flux collected by. the traps, a value considerably less than the 50% previously by Berner and in the above discussion. The authors state that 12% represents a minimum value for the of the Caco3 flux, due to the likelihood of preferential aragonite dissolution to calcite, and the occurrence of shallow sediments with of Another sediment investigation (Honjo, 1 978) found small amounts of aragonite in a trap placed at 5,367 meters in the Sargasso Sea for 75 days. The percentage of calcite in the Caco3 fraction was 93%, with the remai ning 7%. As the pteropod shell showed "evidence of severe dissolution", and this trap was located in waters that are substantially > 30, based on an examination of 45), it is likely that the percentage of actually reaching the trap was much and that a amount of aragonite was dissolved the 75-day sampling period. The sediment investigations in conjunction with the dissolution which is in this

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187 attempted to quantify the calcite and fluxes in the Pacific Ocean et al. 1984). Although the deep deployed in substantially losses due to dissolution minimized by deployment times. The deployed at 100 found an average flux of 30 mg/m2 /day, which is equivalent to 1 1 mg/ cm2 The deployed at 2200 found an of 0.9 mg/m2 /day, which is a flux of 0.03 Note that model of the alkalinity a 2 of mg/cm of Caco3 and extension of Caco3 budget model 0.9 Both values quite close to the flux by the As noted shells and that still visibly motile excluded the flux calculations in the shallow and (1986) used to sediment to question the cited above. and noted that the main of to an is to become "negatively buoyant and often swim This would lead to a substantial enhancement of the flux in due to and substantiated suspicions by of in nets. The state that the flux of in the of et al. ( 1 984) is not to be the of alkalinity i n the Pacific. Two points noting the and study. "sediment (actually a

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188 plankton net) was deployed at a depth of 25 meters. Despite the fact that many of the shells may have entered the trap alive, the natural death of some organisms will have contributed a portion of the shells collected by the trap. Since the number of organisms living "above" the 1 00-meter trap is at least four times the number living "above" the 25-meter trap, this larger population will also contribute a larger number of deceased organisms to the flux measured by the 100-meter traps. Harbison and Gilmer also note that Limacina species have maximum abundances in the 50-100 meter depth range. Although it is likely that the incidence of trap encounters with living organisms is greater at these depths, the population of predators will be greater at these depths, increasing the incidence of predator-prey interactions. As predators discard or excrete pteropod shells, these shells are contributed to the actual downward aragonite flux. Furthermore, at the beginning of their paper, Harbison and Gilmer make the following statement: "Sediment traps have been estimated to collect between 3 and 10% of the material produced in the upper parts of the water column that "rains" down in the form of dead plants and animals, fecal pellets, moults, and empty shells (e.g. BERNER and HONJO, 1981)." It is assumed that the authors are referring to sediment traps deployed in deep water, as was the trap used in the investigation of Berner and Honjo. Traps of the "Soutar" design used by Betzer et al. have been shown to collect at least 95% of the sediment flux in the

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189 Santa Basin et al., 1977). If the authors indeed to sediment collections, it would be useful to examine the fluxes determined by the deployed at 2200 As stated above, the flux determined by these is 0.03 If the maximum and minimum estimates of the collecting efficiency of the are used, the fluxes may be multiplied by a minimum of 10 and a maximum factor of 33. Such an yields a minimum value of 0. 3 mg/cm2 the flux, and a maximum value of 1.0 The minimum value is of the eaco3 flux by the models, and the maximum value is essentially equivalent to the flux. The consistency of these is especially as Harbison and indicate that is likely in shallow Although speculative, it may be possible that a natural biological mimics the mechanism. The process would in in essentially the same as so that the shallow of et al. ( 1984) actually an value the flux. The effects of the biological would not be evident at 400 biological are Predation is the best candidate the biological Harbison and Gilmer state that Clione limacina (a euthecosomatous, non-shelled, feeds on species of thecosomatous (shelled) Limacinas in small in the medusae and fish also known to consume amounts of when available. If pteropods the objects of

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190 opportunistic predation, and the shells are not destroyed by the digestive process, the elimination of consumed shells by predators could mimic the overtrapping of pteropods in sediment traps. Overtrapping would cause large numbers of shells to enter the sample cup, and predation would cause large numbers of shells to be delivered to the water column. Both of these processes are dependent on the patchy distribution of oceanic zooplankton, which is particularly evident in oligotrophic waters. Predation of a large number of pteropods in a short period of time would be a relatively rare occurrence. However, the result of such predation would be delivery of a greater amount of aragonite to the water column than the flux which results from the turnover rate of pteropod populations. If predator encounters with patchy clusters of pteropods occur at approximately the same rate as the interaction of a sediment trap with similar pteropod clusters, the two processes would result in the same magnitude of aragonite flux. Short-term deployments, such as those used by Betzer et al., would be more likely to miss the delivery of predated pteropod shells to depth than long-term deployments. Betzer et al. also suggested that vertically migrating predators could also deliver alkalinity "baggage" deeper in the water column, bypassing the 400-meter horizon. Furthermore, even if the shells were destroyed by digestion, the predator would void excreta high in carbonate content. Therefore, despite the probability of pteropod "overtrapping" in shallow surface traps, there is an apparent consistency between the predictions of the Fiadeiro and "Broecker-Berner" models, the measured aragonite fluxes from the shallow traps, and the "adjusted" fluxes in

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191 the deep traps. This consistency is further supported by observations of large abundances of pteropods in net collections, and equal ratios of pteropods to foraminifera in shallow sediments. These observations lead to the conclusion that aragonite's role in the oceanic co2 system has probably been underestimated. This conclusion is based on very limited data and should be regarded as tentative. Since the magnitude of the aragonite flux in the ocean is by no means conclusively determined, as demonstrated by the limited amount of data that was available for this discussion, the necessity more quantification should be obvious. One of the best ways to carry out this quantification would be to perform short-term sediment trap deployments at depths substantially deeper than the euphotic zone but above the saturation horizon for aragonite. Deploying the traps above the saturation horizon would eliminate any uncertainty regarding the loss of aragonite and calcite to dissolution. Short-term deployments also minimize the possibility of aragonite dissolution due to the oxidation of organic matter and subsequent acidification of seawater contained in the sample cup (Gardner et al. 1983) The entire Atlantic Ocean is an area in which this investigation would be possible. The Atlantic also has areas of high, medium, and low productivity. In order to provide basin-wide estimates of the aragonite flux, deployments should be performed in all three areas. Pteropods and their aragonitic shells are apparently not deliberately designed to confuse the results of sediment trap investigations. However, the biological behavior of the organisms, and the chemical lability of their shells, has caused the results of these initial investigations to be less than definitive with regard to the

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192 aragonite flux to the deep ocean. Due to the importance of the oceanic C02 system in the global co2 system, and in light of the evidence presented here, it is clear that the aragonite "question" is one which requires a definitive answer. It is therefore important to design a program of sampling which can provide such an answer. The magnitude of the aragonite flux in the ocean is a significant, but uncertain factor in models of the oceanic co2 system. Further Investigations Although the research discussed in this dissertation has advanced the understanding of the oceanic co2 system, several important questions remain to be answered. A significant question with regard to aragonite is the actual kinetic equation for small pteropod shells, which make up the majority of the species in the ocean, and will be the most active participants in water column dissolution reactions. I have postulated that the rate constant K for small pteropod shells will be greater than the rate constant determined for the large shells us e d in these experiments, but a reliable determination should be made. This determination will require careful sample handling, and will also require the collection of sufficient numbers of small pteropod shells, so that a distinct dissolution signal can be measured. However, a far more widespread and important question must be answered. The dissolution kinetics of calcite and calcitic material have not been addressed by this research. Because of the far greater amount of calcite sediments in the ocean, and because of their

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193 presence at much greater depths in the ocean than aragonite sediments, it is important to determine the fundamental relationships between calcite dissolution rates and the absolute saturation index for calcite, Models of the oceanic 002 system should incorporate C accurate descriptions of calcite dissolution in the oceans to enable accurate predictions and mass balances. Another factor bearing on calcite sediments is the passive nature of dissolution on the sea floor. Our experimental apparatus mimicked the settling behavior of pteropod shells by causing the shells to fall back and forth within the sample chamber. However, dissolution on the sea-floor will occur in a different manner. Partial saturation may occur directly above the benthic boundary layer, caused by the dissolution of surface eaco3 and slow diffusive transport out of the layer. Also, in this region the more rapid dissolution of aragonite has been theorized to allow partial preservation of calcite in some deep-sea sediments (Morse et al., 1979). Accurate investigation of the sedimentary dissolution process should simulate the environment as closely as possible, and use samples taken from nature when available. Morse (1978) and Keir (1982) have investigated several aspects of sedimentary dissolution kinetics, but new models based on the quantity should be considered. In general, use of the parameter in the models in place of the quantity (1-n) will simply require adjustment of the calculated constants. The question of the magnitude of the aragonite flux in the oceans needs to be resolved. Previously in this discussion, suggestions for the best way to conduct such sediment trap research were made. The primary points presented were that the investigation should take place

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below depths where biological behavior is likely to affect the trapping process, but above depths where chemical reactivity will affect the trap contents. Alternatively, a collection cup could be designed which contained a solution supersaturated with respect to aragonite and calcite, to limit losses due to dissolution. Furthermore, the research should investigate the aragonite flux in areas of low, medi urn and high productivity, and not at tempt to estimate entire oceanic Caco3 fluxes from data obtained at a single location. It is important to emphasize that the oceanographic community requires reliable aragonite flux data. 194 Another aspect of the aragonite flux is that pteropod shells have been theorized to act as vectors of transport for adsorbed metal ions, notably actinides and lanthanides (Shanbhag and Morse, 1982; Thompson, 1987). Adsorption in the surface ocean followed by dissolution and the subsequent release of adsorbed ions at depth can contribute trace metals to deep ocean waters. The goal of the investigations described above is to improve our understanding of the oceanic co2 The research presented in this dissertation has demonstrated a fundamental relationship which governs the dissolution of aragonitic caco3 in the oceans. It should be clear that significant questions remain before a reliable model of the oceanic co 2 system is complete. Whitfield (1984) has suggested that the dissolution of aragonite in shallow regions of the ocean may buffer and slow the effect of increasing atmospheric co2 on the oceans. Because there are uncertainties with regard to the magnitude of the aragonite flux and the nature of the entire dissolution process in the oceans, the process of oceanic neutralization of injected co2

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195 cannot be modeled with complete certainty at the present time. Concluding Remarks This research has provided a new insight regarding the multiple relationships which govern the balance of biological, chemical and geological processes affecting co 2 and caco3 in the It is hoped that this model of the dissolution process will allow a more complete understanding of the dynamic s of the oceanic carbon dioxide system. It is the goal of models to simulate nature, so that past, present, and future trends may be determined. Therefore, the relationships described in this work should enable models to more closely resemble the reality of the natural oceanic environment. In essence, the goal of science should be to understand nature as it is, and this research has advanced the understanding of a natural process. However, this particular natural process is also involved in one of the most significant changes to the natural environment wrought by mankind. The injection of fossil fuel co 2 into the atmosphere / ocean system has significant implications for changes in the global and oceanic environment. The process described here will enable a better understanding of the possible effects of this co 2 input. An important aspect of this research is the fact that the processes described here are part of a continuous cycle. Every day, pteropods live, feed, migrate, reproduce, form their transparent Caco3 shells, and die. Every hour, these shells are sinking towards the bottom of the ocean, into regions of the ocean where they will

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196 dissolve, or come to rest in a sedimentary deposit. If they dissolve, the constituents of their shell are recycled by the circulation of the oceans. It is easy to miss the ubiquitous and ongoing nature of these processes. The paper may state that the dissolution process was demonstrated in an experiment, but the process has been operating long before the experiment was performed, and would continue whether or not it was described in a scientific manner. Our experiments and sampling are only approximations of physical reality. Science attempts. to estimate and quantify, with the goal of describing the past, understanding the present, and predicting the future. However, the operation of the cycle of which aragonite dissolution is a part will continue whether or not scientific descriptions of the cycle are accurate. As science seeks to specify and quantify, the remarkable interrelationships of the natural system should not be overlooked.

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197 LIST OF REFERENCES Acker, J.G, R.H. Byrne, S. Ben-Yaakov, R.A. Feely, and P.R. Betzer (1987) The effect of pressure on aragonitedissolution rates in seawater. Geochim. Cosmochim. Acta 51, p. 2171-2175. Adelseck, C. G. and W. H. Berger ( 1975) On the dissolution of planktonic foraminifera and associated microfossils during settling and on the sea floor. Cushman Found. for Foraminiferal Res., Special Pub. No. 13, pp. 70-81. ---. Almgren, T., D. Dyrssen, and M. Strandberg (1977) Computerized high-precision titrations of some major constituents of seawater on board the R.V. Dmitry Mendeleev. Deep-Sea Res. 24, p. 345-364. Bates, R.G. (1973) Determination of Theory and practice. (Seconded.) John Wiley and Sons, New York,lf79 p. Bates, R.G., R.W. Ramette, and C.H. Culberson (1977) Acid-Base Properties of Tris(hgdroxymethyl)aminomethane (TRIS) Buffers in Seawater from 5 to 40 C. Anal. Chern. 49, p. 867-870. Be', A.W.H. (1977) An Ecological, Zoogeographic, and Taxonomic Review of Recent Planktonic Foraminifera. In: Oceanic Micropaleontology, Vol. 1, (Ed. A.T.S. Ramsay), Academic Press, New York, pp. 1-100. Be', A.W.H., and R.W. Gilmer (1977) A zoogeographic and taxonomic review of Euthecosomatous In: Oceanic Vol. 1, (Ed. A.T.S. Ramsay), Academ1c Press, New York, pp. 733-80 Ben-Yaakov, S., and I.R. Kaplan (1971) Deep-sea in-situ calcium carbonate saturometry. J. Geophys. Res. 76, pp. 722-731. Ben-Yaakov, S., E. Ruth, and I.R. Kaplan (1974) Calcium carbonate saturation in the NE Pacific: in-situ determination and geochemical implications. Deep-Sea Res. 21, pp. 229-243. . Berger, W.H. (1967) Foraminiferal ooze: solution at depths. Science 156, pp. "383-385. Berger, W.H. (1968) Planktonic foraminifera: selective solution and paleoclimatic interpretation. Deep-Sea Res. 15, pp. 31-43. Berger, W.H. (1975) Deep-sea carbonates: Dissolution profiles from foraminiferal preservation. Cushman Found. for Foraminiferal Res., Special Pub. No. 13, pp. 82-86.

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198 Berger, W.H. (1978 ) Deep-sea carbonates Pteropod distributi on and the aragonite compensation depth. Deep-Sea Res. 25, pp. 447-452. Berner, R. A. ( 1971 ) Principles of Chemical Sedimentology, McGraw-Hill, St". Louis, 240 p. Berner, R.A. (1976) The solubility of calcite and aragonite in seawater at atmospheric pressure and 34.5/00 salinity. Am. J. Sci. 276, p. 713-730. . --Berner, R.A. (1977) Sedimentation and dissolution of pteropods in the ocean. In: The Fate of Fossil Fuel co2 in the Oceans, Eds. N. R. Andersen and A. Malahoff, Plenum Press, New York, pp. 243-259. Berner, R. A. and J. W. Morse ( 1 97 4) Dissolution kinetics of calcium c arbona t e in seawater. IV. Theory of calcite dissolution. Am. Sci. 274, pp. 108-134. . Berner, R.A., E.K. Berner, and R.S. Keir (1976) Aragonite dissolution on the Bermuda Pedestal: its depth and geochemical significance. Earth Planet. Sci. Lett. 30, pp. 169-178. . Berner, R.A. and S. Honjo (1981) Pelagic sedimentation of aragonite; Its Geochemical significance. Science 211, pp. 940-942. Betzer, P.R., W.J. Showers, E.A. Laws, C.D. Winn, G.R. DiTullio, and P.M. Kroopnick ( l984a) Primarg produ ctivity and particle fluxes on a transect of the equator at 153 W in the Pacific Ocean. Deep-Sea Res. 31, pp. 1-11. Betzer, P.R., R.H. Byrne, J.G. A cker, C.S. Lewis, R.R. Jolley, and R. A. Feely (l984b) The Oceantc Carbonate System: A Reassessment of Biogenic Controls. Science 226, pp. 1 0 74-1077. Biscaye, P.E., V Kolla, and K.K. Turekian (1976 ) Distribution of calcium carbonate in surface sediments of the Atlantic Ocean. Jour. Geophys. Res. 81, pp. 2595-2603. Boudreau, B. P. and N L. Gu inasso ( 1 982) The influence of a diffusive sublayer on accretion, dissolution, and diagenesis at the sea floor. In: The Dynamic Environment of the Ocean Floor (Ed. K.A. Fanning and F.T. Manheim), Lexington Books, Lexington MA, 502 p. Bradshaw, A.L., P.G. Brewer, D.K. Shafer, and R.T. Williams (1981) Measurements of total carbon dioxide and alkalinity by p otentiometric titration in the GEOSECS program. Earth. Planet. Sci. Lett. 55, pp. 99-115. Bramlette, M.N.A. ( 1961) Pelagic Sediments. In: Oceanography, Ed. M. Sears, AAAS Publication, Washington, D.C., pp. 345-366. Broecker, w.s. (1971) Calcite accumulation rates and glacial to interglacial changes in oceanic mixings. In: Late Cenozoic Glacial Ages, Ed. K.K. Turekian, Yale Univ. Press, New Haven, pp. 239-265.

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199 Broecker, W.S. (1974) Chemical Oceanography. 214 p., Harcourt Brace Jovanovich; Inc. NY. Broecker, W.S. and T.S. Peng (1982) Tracers in the Sea. 690 p., Lamont-Doherty Geological Observatory Pub., New York.--. Broecker, W.S. and T. Takahashi (1977) Neutralization of fossil fuel C02 by marine calcium carbonate. In: The Fate of Fossil Fuel co2 in the oceans, Eds. N.R. Andersen and A. Malahoff, Plenum Press, NewYork, pp. 213-241. Burton, W.K. and N. Cabrera (1949) Crystal growth and surface structure, Part I. Faraday Soc. Disc. 5, pp. 33-39. Burton, W.K., N. Cabrera, and F.C. Frank (1951) The growth of crystals and the equilibrium structure of their surfaces. Royal Soc. London Phil. Trans. A-243, pp. 299-258. Byrne, R.H. (1984) Absorbance corrections in self-adjusting, variable path-length-diameter, high pressure cells. Rev. Sci. Instrum. 55 (1) t pp. 131-132. --. --. Byrne, R.H., J.G. Acker, P.R. Betzer, R.A. Feely, and M.H. Cates ( 1984) Water column d issolution of aragonite in the Pacific Ocean. Nature 312 (5992), pp. 321-326. Byrne, R.H., G. Rebert-Baldo, S.W. Thompson, and C.T.A. Chen (1988) Seawater pH measurements: An at-sea comparison of potentiometric and spectrophotometric methods. Deep-Sea Res., in press. Chen, C. (1964) Pteropod ooze from the Bermuda Pedestal. Science 144, pp. 60-62. Chou, L. (1987) Dissolution kinetics of carbonate minerals. PRF annual report, 34 pages Cooke, R.C. and P.E. Kepkay (1980) Solubility of aragonite in seawaterII: Effect ofpressure on the onset and maintenance of dissolution. Geochim. Cosmochim. Acta. 44, pp. 1077-1080. Culberson, C.H. (1972) distribution of carbon dioxide. 178. Processes affecting the oceanic Ph.D. thesis, Oregon State Univ., p. Culberson, C .H., D.R. Kester, and R .M. Pytkowicz (1967) High Pressure Dissolution of Carbonic and Boric Acids in Seawater Science 157, pp. 59-61. Culberson, C .H. and R.M. Pytkowicz (1968) Effect of pressure on carbonic acid, boric acid, and the pH in seawater. Limnol. Oceanogr. 13, pp. 403-407.

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Culkin, F. (1965) The major constituents of seawater. In: Chemical Oceanography, Vol. 1 Eds. J. P. Riley and G. Sk irrow Academic Press, New York, pp. 121-158. Culkin, F. and R.A. Cox (1966) Sodium, potassium, magnesium, calcium, and strontium in seawater. Deep-Sea Res. 13, pp. 789-804. 200 Deuser, W.G., E.A. Ross, and W.F. Anderson (1981) Seasonality in the supply of sediment to the deep Sargasso Sea and-implications for the rapid transfer of matter to the deep sea. Deep-Sea Res., 28, pp. 495-505. --. Edmond, J.M. and J.M.T.M. Gieskes (1970) On the calculation of the degree of saturation o-f seawater with respect to calcium carbonate under in-situ conditions. Geochim. Cosmochim. Acta 34, pp. 1261-1291. Feely, R.A., R.H. Byrne, P.R. Betzer, J.F. Gendron, and J.G. Acker (1984) Factors influencing the degree of saturation of the surface and intermediate waters of the North Pacific Ocean with respect to aragonite. Jour. Geophys. Res. 89 (C6), pp. 10,631-10,640. Fiadeiro, M. (1980) The alkalinity of the deep Pacific. Earth Planet. Sci. Lett. 49, pp. 499-505. Gardner, W.D., K.R. Hinga, and J. Marra ( 1983) Observations on the degradation of biogenic material in the deep ocean with implications on accura c y of sediment trap fluxes. Deep-Sea Res. 41, pp. 195-214. Hansson, I. (1972) An analytical approach to the carbonate system in seawater. thesis, Univ. of Goteborg, Sweden. Hansson, I. (1973) A new set of pH-scales and standard buffers for seawater. Deep-Sea Res. 20, pp. 479-491. Harbison, G.R. and R.W. Gilmer (1986) Effects of animal behavior on sediment trap collection-s: implications for the calculation of aragonite fluxes. Deep-Sea Res. 33, pp. 1017-1024. . Hawley, J.E. and R.M. Pytkowicz (1969) Jolubility of calcium carbonate in seawater at high pressures and 2 C. Geochim. Cosmochim. Acta 33, pp. 1557-1561. Honjo, S. (1976) Coccoliths: Production, transportation, and sedimentation.Mar. Micropaleo. 1, pp. 65-79. Honjo, s. (1977) Biogenic carbonate particles in the ocean; Do they dissolveinthe water column? In: The Fate of Fossil Fuel co2 in the Oceans, Eds. N.R. Andersen and A. Malahoff, Plenum Press, New 269-294 Honjo, s. (1978) Sedimentation of materials in the Sargasso Sea at a 5,367 m deep station. J. Mar. Res. 36, pp. 489-492.

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201 Honjo, S. (1980) Material fluxes and modes of sedimentation in the mesopelagic and bathypelagic zones. Jour. Mar. Res. 38, pp. 53-97. Honjo, S. and J. Erez (1978) Dissolution rates of calcium carbonate in the deep ocean: An in-situ experiment in the North Atlantic Ocean. Earth Planet. Sci. Lett. 40, pp. 287-300. . Honjo, S. and M.R. Roman (1978) Marine copepod fecal pellets; production, preservatton, and sedimentation. Mar. Res. 36, pp. 45-57. Honjo, S., S.J. Manganini, and J.J. Cole (1982) Sedimentation of biogenic matter in the deep ocean. Res. 29(5A), pp. 609-625. Kobayashi, H.A. (1974) Growth cycle and the related vertical distribution of the thecosomatous pteropod Spiratella ( "Limacina") helicina in the central Arctic Ocean. Mar. Biol. 26, pp. 295-301. Ingle, S.E. (1975) Solubility of calcite in the ocean. Mar. Chern. 3, pp. 301-319. --. Ingle, S.E., C.H. Culberson, J.E. Hawley, and R.M. Pytkowicz (1973) The solubility of calcite in seawater at atmospheric pressure and 35 00 salinity. Mar. Chern. 1, pp. 295-307. . Keir, R.S. (1980) The dissolution kinetics of biogenic calcium carbonates in seawater. Geochim. Cosmochim. Acta 44, pp. 241-252. Keir, R. S. ( 1982) Dissolution of calcite in the deep-sea: Theoretical prediction for the case of uniform size particles settling into a well-mixed sediment. Am. 282, pp. 193-236 leNoble, W.J. and R. Schlott (1976) All-quartz optical cell of constant diameter for use in high-pressure studies. Rev. Sci. Instrum. 47, pp. 770-771. Lyman, J. (1957) Buffer mechanism of seawater. Ph. D. Thesis, University of california at Los Angeles, 196 p. Macintyre, W.G. (1965) The Temperature Variation of the Solubility Product of Calcium Carbonate in Seawater. Manuscript, Rep. Ser. No. 200, Fish Res. Board, Canada, 153 pp. Mackenzie, F. T., W.O. Bischoff, F.C. Bishop, M. Loijens, J. Schoonmaker and R. Wollast. (1983) Magnesian calcites: Low-temperature occurrence, solubility, and solid solution behavior. In: Carbonates: Mineralogy and Chemistry. Ed. R.J. Reeder. Mineral Soc. Amer., Reviews in Mineralogy 11, pp; 97-144. McGowan, J.A. (1960) The systematics, distribution, and abundance of the Euthecosomata of the North Pacific. PhD Dissertation, University of California, San Diego, 197 pp.

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202 McGowan, J.A. (1963) Geographical variation of Limacina helicina in the North Pacific. Systematics Assoc. Pub. No. 5, Speciation in the Sea, pp. 109-128.Mehrbach, C., C.H. Culberson, J.E. Hawley, and R.M. Pytkowicz (1973) Measurements of the apparent dissociation constants of carbonic accid in seawatr at atmospheric pressure. Limnol. Oceanogr. 18, pp. 897-907. Melgue, M. and J. Thiede (1974) Facies distribution and dissolution depths of surface sediment components from the Vema and the Rio Grande Rise. Mar. Geol. 17, pp. 341-353. Millero, F.J. (1979) The thermodynamics of the carbonate system in seawater. Geochim. Cosmochim. Acta 43, pp. 1651-1661. . Milliman, J.D. (1977) Dissolution of calcium carbonate in the Sargasso Sea (Northeast Atlantic). In: The Fate of Fossil Fuel in the Oceans, Eds. N.R. Andersen and A. Malahoff, Plenum Press, ew York, pp. 541-654. Morse, J.W. (1974a) Dissolution kinetics of calcium carbonate in seawater. III. "A new method for the study of carbonate reaction Am.274, pp. 97-107. . Morse, J.W. (1974b) Dissolution Kinetics of calcium carbonate in seawater. V. Effects of natural inhibitors and the position of the chemical lysocline. Am. 274, pp. 638-647. . Morse, J.W. (1978) Dissolution kinetics of calcium carbonate in seawater. VI. The near-equilibrium dissolution kinetics of calcium carbonate in deep-sea sediments. Am. Sci. 278, pp. 344-353. Morse, J.W. (1983) The kinetics of calcium carbonate dissolution and precipitation. In: Carbonates: Mineralogy and Chemistry. Ed. R.J. Reeder. Mineral Soc. Amer., Reviewsin Mineralogy 11, pp. 227-264. Morse, J.W. and R.A. Berner (1979) The chemistry of calcium carbonate in the deep oceans. In: Chemical Modeling in Aqueous Systems: Speciation, Sorption, Solubility, and Kinetics. Ed. E. Jenne, ACS Symposium Series 93, American Chemical Society, Washington; D.C., pp. 499-536. Morse, J.W., J. deKanel, and K. Harris (1979) Dissolution kinetics of calcium carbonate in seawater. VII. The dissolution kinetics of synthetic aragonite and pteropod tests. Am. 279, pp. 488-502. . Mucci, A. (1983) The solubility of calcite and aragonite in seawater at various salinities, temperatures, and one atmosphere total pressure. Am. 283, pp. 780-799. . Peterson, M.N.A. (1966) Calcite: rates of dissolution in a

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203 vertical profile in the central Pacific. Science 154, pp. 1542-1544 Plath, D.C. ( 1 979) The solubility of CaCO in seawater and the activity coefficients in solutions. M.S. Thesis Oregon State Univ. Corvallis. 92 p. Plath, D C and R.M. Pytkowicz (1980) The solubility of aragonite at 25C and 32.62 /00 salinity. Mar. Chern. 10, p. 3-7. Plunmer, L.N., T .M.L. Wigley, and D.L. Parkhurst (19'b8) The kinetics of calcite dissolution in C02._-water systems at 5 to 60 and 0 to 1 atm co2 Sci. z r8, pp. 179 -2 1 6. Plummer, L .N., D.L. Parkhurst and T.M.L. Wigley (1979) Critical review of the kinetics of calcite dissolution and precipitation. In: Chemical Modeling in Aqueous Systems: Speciation, Sorption, Solubility, and Kinetics. Ed. E Jenne ACS Symposium Series 93, Arnedcan Chemical Society Washington, D C pp. 537573. Price, B.A. J.S. Killingley, and W.H. Berger (1985) On the pteropod pavement of the eastern Rio Grande Rise. Mar. Geol 6 4 pp. 217-235 . --Rebert Baldo, G M.J. Morris, and R .H. Byrne (1985) Spectrophotometric determination of seawater pH using phenol red. Anal. Chern. 57, pp. 2564-2567. Sass E J.W. Morse, and F.J. Millero (1983) Dependence of the value s of calcite and aragonite thermodynamic solubility products on ionic models Am. 238, pp. 2 18-229 Sayles, F.L. (1980) The solubility of CaCO in seawater at 2C based upon in-situ sampled pore water Mar. Chern. 9, pp. 223-225 . Sayles, F.L (1985) CaCO solubility in marine sediments : Evidence for equtlibrium and n%n-equilibriu.rn behavior. Geochim. Cosmochim. Acta 49, pp. 877-888 Schink, D.R. and N .L. Guinasso, Jr. (19 77) Modelling the Influence of Bioturbation and other processes o n calcium carbonate dissolution at the sea floor. In: The Fate of Fossil Fuel co2 in the Oceans Eds. N .R. Andersen and A. Malahoff, Plenum Press, New York, pp. 375-400; Schindler, P. ( 1967) Heterogenous Equilibria Involving Oxides, Hydroxides, Carbonate and Hydroxide Carbonates. In: Equilibrium Concepts in Natural Water Systems Ed. W. Stumm, Advances in Chemistry Series No. 67, ACS, Washington D C pp. 196-221. Schindler, P., H. Althaus, F. Hofer, and W Minder (1965) Loslichkeitsprodukte v o n Metalloxiden und hydroxiden: Loslichkeitsprodukte von Zinkoxid, Kupferhydroxid und Kupferoxid in Abhangigkeit von Teilchengrosse und molarer Oberflache Ein Beitung

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205 APPENDICES

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206 APPENDIX 1: Calculations of K' as a function of Temperature, spa Salinity, and Pressure. The calculation of the saturation state of the oceanic water column (or in our pressurized experiments) is based on a series of equations. These equations have been determined for elements of the seawater alkalinity system and the carbonate system. Equations are used to calculate the value of the following constants: a) the first and second apparent dissociation constants of carbonic acid in seawater CK1 and K 2'); b) the first apparent dissociation constant of boric acid in seawater (K8'); c) the values of the apparent solubility products for calcite and aragonite (K'spc and K'spa) Several determinations of carbonate system constants as a function of temperature, salinity, and pressure have been performed. As a result, there are several available 'sets' of equations available which will provide descriptions of the oceanic carbonate system. Judicious investigators of this system must choose a set of equations which are based on reliable data, and which will provide the best description of actual oceanic conditions. Furthermore, the solubility products of calcite and aragonite also vary as a function of temperature, salinity, and pressure, and equations are necessary to describe their functional dependence. This situation has been complicated by the variation in the measured solubility of aragonite, which has been shown to decrease as equilibration time increases (Morse et al., 1980).

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APPENDIX 1, continued 207 This appendix will not attempt to survey all of the equations that have been formulated to describe the oceanic carbonate system. Equations pertinent to the dissociation constants of carbonic acid may be found in Edmond and Gieskes ( 1970), Culberson et al. ( 1972), Mehrbach (1973), Millero (1979), Takahashi et al. (1980) and Feely et al. (1984) Lyman (1957) and Hansson (1972) examined the functional dependence of K8', and their data was re-analyzed by Millero (1979). As research into the oceanic carbonate system has progressed, each research group has formed its own set of equations to describe the system. As examples, examine Takahashi et al. (1980), which describes the calculations used by the GEOSECS expeditions, and Feely et al. (1984), which describes the equations utilized in his research. The research described in this dissertation has conformed to the equations used by Feely et al. (1984), with regard to the calculation of the dissociation constants. In this appendix, the equations used to describe the solubilty products of calcite and aragonite as functions of temperature, salinity, and pressure will be examined. Also, the results of several determinations of K'sp for both minerals will be delineated. The various values of 6V, the partial molal volume change for calcite and aragonite dissolution, will be also be delineated. Small variations in the value of 6V profoundly influence the calculation of the saturation state under pressure. This appendix should help to alleviate some of the confusion caused by the various values and equations that have been published in the scientific iterature. Several determinations of the solubility product of calcite and aragonite have been performed. Table 7 lists values of K' and K' spc spa

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APPENDIX 1 continued 208 TABLE 7 : Values of K and K' at 25 degrees Centigrade and spa spa 35/00 salinity. The studies which determined these values are also shown. The units for all values are (mol/kg) 2 Study K' spa K spa Macintyre, 1965 4 .57 X 107 7,33 X 107 Ingle, 1975 4.6 X 1 0_7 Berner, 1976 (calc.) 5 .49 X 107 8.11 X 107 Plath & Pytkowicz, 1980 4.7 X 107 8.69 X 107 Morse et al., 1980 4.39 X 107 6.65 X 107

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APPENDIX 1 continued 209 determined at 35 parts per thousand salinity and 25 degrees Centigrade. For calcite, Berner's calculated value has generally been abandoned. However, his value influenced research for several years, including Keir (1980). Also, Berner's ratio of the aragonite and calcite solubility products, determined theoretically, has been used by several investigators, notably the GEOSECS expeditions (Takahashi et al. 1980) If Berner's value is not included, the values of the calcite solubility product vary only by a factor of 1.07. However, note that the values of the aragonite solubility product vary by a factor of 1.3. Figure 14, which was shown previously in the dissertation (Byrne et al., 1984, adapted from Morse et al., 1980), demonstrates that the determined value of the solubility product varies as a function of the equilibration time. The value of the aragonite solubility product from Morse et al. (1980) and Mucci (1983), 6.65 x 107 mo12 kg2 is now generally accepted. Mucci's determined K'spa for pteropod shells is only slightly higher, 6.73 x 10-7 mo12 kg2 Keir (1980) uses a value of 0.50 (mmol/1) 2 for the solubility product of calcite. Conversion of this value to (mol /kg)2 gives 4.774 x 10-7, which Keir states is "very close" to the value of Ingle (1973), namely 4.6 x 107 (mol/kg)2 Table 8 shows values of K'spc determined between 2 and 5 C, for 35 /00 and 1 atmosphere pressure. The values from Sayles (1985) were determined by pore water saturometry. Sayles (1980) gives a value of K' equal to 4.61 x 107 (mol/ 1) 2 Sayles (1980) appears to state spc . that Ingle's value of K'spc is 4.8 x 107 (mol/1) 2 but this value is actually for (mol/kg)2 units. Sayles' 1985 values, which are for

PAGE 236

APPENDIX 1 continued 210 TABLE 8 : K spc values obtained between 2 and 5C, at 35/00 and 1 atmosphere pressure. Study Ingl e 1975 (from f i gure) 4 .765 X 107 (ool/kg) 2 Ingle (cited by Sayles 1985) 4.85 X 107 (mol/kg) 2 Sayles, 1980 4.61 X 107 (mol/kg) 2 Sayles, 1985 4.95 X 107 (mol/kg) 2

PAGE 237

APPENDIX 1, continued (mol/kg)2 units, are shown in Table 8. 211 The original calculation algorithm used in this study followed the GEOSECS expedition algorithm, in which the solubility product of calcite was calculated according to the relationship determined by Ingle et al. ( 1973), and the solubility product of aragonite was calculated by multiplying by a constant ratio, K' IK' This spa spc point is relevant to the following discussion. Several investigators have analyzed the temperature and salinity dependence of the caco3 solubility products. Edmond and Gieskes ( 1970) used the following equation for the solubility product of calcite: K'spc ( 0.1614 + 0.02892 Cl .0063 t ) x 106 where Cl is the chlorini ty, and t is the temperature in degrees Centigrade. The only difference in the equation for K'spa is the value of the first term, 0.5115 instead of 0.1614. Ingle (1973) formulated the following expression for the solubility product of calcite: K' (1 atm) = [ -34.452-39.866 S112 + 110.21 logS spc 7.5752 X 106 T2 ] X 107 Millero ( 1979) fitted Ingle's data to equations of a different form for the solubility products of calcite and aragonite. Data from Ingle (1973) was used for calcite; data from Berner (1976) was used for aragonite. The calculation of either solubility product involves

PAGE 238

APPENDIX 1, continued 212 the influences of temperature and salinity. The values calculated by the following expressions are designated K for calcite, and K for c a aragonite. Note that the only difference in these two expressions is the value of the first term. ln Kc = 303.130813348.09/T-48.7537 ln T ln Ka 303.5363 -13348.09/T-48.7537 ln T The salinity dependence of either solubility product is determined with the following expression (calcite is used as the example): ln K' -ln K (1.6233-118.64/T) S112-(0.07 S) c c The expression used by Feely et al. ( 1984) to calculate the solubility product of aragonite was obtained from Mucci (preprint of 1983 ms.). Although the equation used by Feely et al. is not identical to the equation appearing in Mucci (1983), the two equations give identical results. The expression for the salinity dependence is the same: K* (-.068393 + .0017276 T + 88.135/T) S112 + (-0.10018 S) + 0.0059415 S1 5 Feely et al. 's (1984) equation: 0 log Ksp 505.163 + 0 .115304 T-15169.2/T200.841 log T

PAGE 239

APPENDIX 1 continued 213 Mucci (1983) equation: log Kspo -171.945-0.077993 T + 2903.293/T + 71.595 log T In both cases, the apparent solubility product is then calculated: log K' K + log K 0 sp sp The expression which was used in this dissertation to calculate the solubility product of aragonite is of the same form as Ingle's equation for calcite: K' spa [-49.842-57.68 S1 / 3 + 159.42 logS 1.096 X 105 r2J 10-7 In the course of describing the calculation algorithm, the difference between our algorithm and that of Feely became apparent. However, water column data from Feely et al. (1984) had previously been used to test the computer program, and the difference between the calculated values of n was small. (The equations used to calculate the effect of pressure will be described below). The reason for the slight difference between the two algorithms is the actual K' value returned by the two equations. Mucci (1983) spa reports a value of 6.65 x 107 (mol/kg) 2 for conditions of 25C and 35 0 /00 salinity. However, his equation returns a value of 6.48 x 107 (mol/kg) 2 for the same conditions. The modified equation of Ingle

PAGE 240

APPENDIX 1 continued 214 used in our method returns a value of 6.68 x 107 (mol/kg) 2 for the standard conditions of 25C and 35/00 Mucci's equation was deternined in a salinity range of 5 to 44 I 0 0 and Ingle's was determined in a salinity range of 27 to 43/00 As seawater salinity values range from 34 to 37/00, Ingle's equation may yield more accurate values in this range. Calculation of the water column saturation state at depth also involves the influence of pressure. Several investigators have described the effect of pressure on the oceanic carbonate system. As this calculation is vitally involved with many of the results in this dis.sertation, it is necessary to examine these calculation methods. Edmond and Gieskes (1970) used the following relationship for the effect of pressure on the solubility product of calcite or aragonite: ln (K ') /(K ') = (P-1)/RT sp P sp 1 In this expression, the subscripts P and 1 designate pressurized and atmospheric pressure conditions. R is the gas constant (units of cm3atm / Kmole), Tis the Kelvin temperature, and the partial molal volume change for the dissolution of calc i urn carbonate, in cm3/rnole units. The expression was based on the data of Hawley and Pytkowicz (1969). Edmond and Gieskes used the following expressions for the calculation of (35.4 0.23 t) c (32 8 o.23 t) a In these expressions, th t t 1n C. t 1s e empera ure

PAGE 241

APPENDIX 1, continued 215 The GEOSECS expedition (Takahashi et al., 1980) on the of the oceans, uses a formulated by (1972): K' = K' 0.20 t)CP spc,P spc, 1 K' = K' e(33.3 t)CP spa, P spa, 1 CP (P-1)/83.143 T In this P is the in T is the in K, and t the in c. Ingle (1973) deterninations of calcite and solubility She fit the to an equation of the following foro: -6V(P-1) + 0.5 6K(P-1) 2 6K is the compressibility the defined as (1979) used a for the influence of In equation, the (P-1) is replaced simply by P. in Mucci et al. (1982), an which did not include the (6K) was used. This is the which was used in the of this form: ln (K' )/(K' ) sp,P sp,1 (6V/RT)P

PAGE 242

APPENDIX 1, continued 216 Feely et al. (1984) also used this equation for the effect of pressure. It is clear that the use of the above equations requires values of for the dissolution of aragonite and calcite. For simplicity, the examination will be restricted to low-temperature determinations. Table 9 shows the values of V which have been determined by the researchers shown (or calculated from their equations), all at 2c. Both Morse and Berner (1978) and Millero (1979) indicate that the difference in the solid phase molar volumes of calcite and aragonite is 2.8 cm3 / mole. On a theoretical basis, this difference should be the exact difference between the values of for calcite and aragonite dissolution. Note that the difference of the "calculated" values of Edmond and Gieskes, and Culberson, are close to this value. However, the best experimentally-determined values of for calcite dissolution (Ingle, 1973; Sayles, 1980, 1985) are over 10 cm3 /mole greater than experimentally-determined values for aragonite dissolution. It is intriguing that the value calculated in this study, -37.0 cm3 /mole, lies midway between the value from laboratory determinations and the value determined from the molar volume difference of the solid phases. Also, our other analyses appeared to indicate that the use of a higher value for the value of is justified. The "best" value of for the dissolution of aragonite is difficult to determine. Mucci et al. (1982) states that the average value of Hawley and Pytkowi cz, Cooke and Kepkay, and Ingle is -31.3 cm3/mole. This value was adopted by Feely et al. (1984). By averaging

PAGE 243

APPENDIX 1 continued 217 TABLE 9: Values of 6V for aragonite and calcite dissolution, all at 2C. The studies from which these values were obtained are also shown. Units are cm3 / mol. Study Hawley & Pytkowicz, 1969 Ingle, 1975 Cooke and Kepkay, 1980 Sayles (at 3 1.5C) 1980 1985 Morse et al., 1980 Edmond & Gieskes, 1970 (calc.) C ulberson, 1972 ( calc.) Acker et al., 1987 This dissertation 33.1 31.8 42.3 29.0 43.8 41.2 31.3 32.34 34.94 3 2.86 35.4 36.5 37.0

PAGE 244

APPENDIX 1 continued 218 all of the laboratory determinations (including our own) the "calculated" values, and also believing higher values of 6.V are justified, an estimate of -32.6 cm3/mole is obtained. Note, however, that this value of 6.V was not used to obtain the kinetic results of this study. It should be emphasized that further investigation of the solubility of calcite and aragonite under pressure would be useful. The discrepancy between laboratory determinations of 6.V for aragonite, and theoretical calculations of 6.V, should be alleviated.

PAGE 245

219 APPENDIX 2 Data Dissolution The column with the heading "File" is the name of the datafile in the Scien c e system. The "E" notation in the "RATE" colwnn the exponent of 10 which is multiplied by the I.e. 2.2 E-7 2.2 x 107 TA and EC02 in meq/kg Dataset #1: Shell mass 33.40 File Temp (K) PSI Salinity TA 1124AT 1124FK 11244K 1125AT 1125TK 1125MK 11262K 11265K 275.16 14.6 II 1555 II 4050 II 14.6 II 3520 II 4550 II 2065 II 5005 36.248 II II II II II II II 2.422 2;439 2:428 2.437 2.482 Dataset #2: Shell mass33.40 File Temp (K) PSI Salinity TA 1127 AT 1127FK 1127TK 11281 K 11283K 1128MK 1129AT 11294K 1129VK 275.16 14.6 II 1570 II 3510 II 1030 3030 4500 II 14.6 4035 5570 36.248 " " " II II 2.432 2.457 2;480 2:453 2;453 2;487 EC02 2.389 2; 391 2;398 2:389 2:394 2;409 2.393 2;400 EC02 2.389 2;392 2. 401 2; 389 2;394 2 ;406 2.392 2 :392 2;409 RATE (meg/min) 2.2 E-7 8;2 E 7 2 ;23 E-6 4:7 E-7 1 .84 E-6 2:55 E-6 7;9 E-7 2:41 E-6 RATE (meg/min) 3.8 E-7 1;06 E-6 2.35 E-6 4:5 E-7 1 ;5 E-6 2 :23 E-6 3.1 E-7 2:03 E-6 3.13 E-6

PAGE 246

APPENDIX 2, continued 220 Dataset #3: Shell mass 33.40 milligrams File PSI Salinity TA tco 2 RATE (meg/min) 1130AT 275.16 14. 6 36.248 2 .430 2.389 5.3 E-7 1130TK 11 3550 11 2:439 2:393 1 :6 7 E-6 11305K II 5030 II 2.467 2:407 3:19 E-6 121YK II 3830 11 2:457 2:393 2.46 E-6 121ZK 11 5175 II 2 .496 2.412 3.36 E-6 112582 II 465 II 2:458 2:389 9 : 0 E-8 1223K II 3080 II 2 .468 2:394 1 .58 E-6 122VK II 5430 II 2:495 2 ;407 2:99 E-6 Dataset 114 : Shell mass 12.502 milligrams File Temp (K) PSI Salinity TA tC02 RATE (meg/min) DIS301 278.16 15.0 34.326 2.424 2 .453 6 6 E-7 DIS302 II 1540 II 2 :439 2 .460 1:36 E-6 DIS311 II 15.0 34.096 2:401 2:401 7:3 E-7 DIS312 II 900 II 2;415 2.408 1:26 E-6 DIS321 II 15. 0 34.326 2:424 2:453 1 :23 E-6 DIS322 II 1480 II 2.462 2.472 1 :1 E-6 DIS33 II 590 33.949 2:353 2:315 4;4 E-7 DIS341 II 15.0 34.568 2:480 2 .465 7.0 E-7 DIS342 II 2930 II 2:49 4 2:472 1:16 E-6 DIS351 II 15.0 34.349 2.431 2.460 9:2 E-7 DIS352 II 1550 2 :450 2:470 6;2 E-7 DIS39 II 1460 II 2; 431 2 :460 1 93 E-6 Last 3 points in Figure 32' ("+" symbols) DIS54 2230 34.506 2.508 2 .502 2.54 E-6 DIS56 590 33:960 2:374 2:338 7:2 E-7 DIS57 II 460 33.84 4 2:353 2:286 7:0 E-7 Dataset /15: Shell mass 45.608 milligrams File Temp (K) PSI Salinity TA tC02 RATE (meg/min) MR15N1 277.66 1540 34.754 2 .334 2 .315 1 .88 E-6 MR15N2 2895 2:350 2:323 2:04 E-6 MR15N3 II 4525 II 2 .368 2:332 3:23 E-6 MR16N1 278.16 1650 34.669 2:391 2:33 1 2:1 E-7 MR16N2 11 2950 II 2:394 2:333 1 .23 E-6 MR16N3 II 4400 II 2:405 2 .338 1:54 E-6 MR17N1 278.56 1360 34.767 2 .383 2 :313 2:0 E-7 MR17N2 II 4000 II 2:386 2;314 9.2 E-7 MR17N3 II 3865 II 2.392 2:317 1 03 E-6 MR17N4 II 5700 II 2:399 2:32 1 2:89 E-6 11L11 points Figure 18 11+11 points Figure 31 JA1318 278.16 4440 34.707 2 .448 2.318 7 5 E-7 JA2318 II 4300 II 2:456 2.322 7;0 E-7 JA319 277.46 4500 2:448 2:318 7 ;O E-7 JA319L 277.56 5170 34.721 2:380 2.268 2.69 E-6

PAGE 247

APPENDIX 2, continued 221 Dataset #6: Shell mass-13.50 milligrams Temp (K) PSI salinity TA File EC02 RATE (meg/min) 1019AT 275.16 14.6 35.081 2.439 2.446 6.6 E-7 10194K II 4040 2;451 2;452 2:3 E-6 10195K II 5000 II 2;493 2.473 2.99 E-6 1020AT II 14.6 2 ;450 2;446 5:2 E-7 1020FK II 1540 II 2;459 2;451 1.4 E-6 10203K II 3080 II 2;484 2;463 1 :99 E-6 1021AT II 14.6 II 2;430 2;446 1 .11 E-6 10212K II 2085 II 2;451 2:457 2:26 E-6 1021MK II 4470 II 2;491 2.477 2:89 E-6 Dataset lf7: Shell mass 13.50 milligrams File Temp (K) PSI salinity TA EC02 RATE (meg/min) 1022AT 275.16 14.6 35.081 2.446 2.446 8.5 E-7 10221 K II 1055 II 2:463 2:454 1:05 E-6 1022TK II 3520 II 2;484 2.464 2:39 E-6 23IAT II 14.6 II 2:440 2:446 1; 19 E-6 23I2K II 2060 II 2:461 2;456 2:38 E-6 23IMK II 4500 II 2:503 2:477 3 :75 E-6 23IIAT II 14.6 II 2.442 2;446 1 ;25 E-6 23II2K II 2050 II 2:464 2:45 7 2:3 E-6 23IIMK II 4510 II 2:505 2:478 3:51 E-6 Dataset #8: S h ell mass 13.50 milligrams File Temp (K) PSI Salinity TA EC02 RATE (meg/min) 1024AT 275.165 14.6 35.081 2 .449 2 .446 1 .01 E-6 10241K 275:16 1050 II 2 :467 2;455 1 ;48 E-6 1024TK II 3490 II 2;493 2 .468 2;68 E-6 1025AT II 14.6 II 2:433 2;446 9:8 E-7 1025QK II 2550 II 2 451 2:455 2 78 E-6 10254K II 4040 II 2;501 2;480 3:59 E-6 1026AT II 14. 6 II 2;446 2;446 1 :27 E-6 102688 II 835 II 2 ;468 2:45 7 1 ;43 E-6 10263K II 3050 II 2 .493 2;470 2;88 E-6 1027AT II 14.6 II 2 ;440 2;4 4 6 1:46 E-6 10273K II 3060 II 2.465 2 .459 3.25 E-6 10275K II 4965 II 2:453 2;495 4;54 E-6

PAGE 248

APPENDIX 2, continued 222 Dataset #9: Shell mass 10.02 milligrams Temp (K) PSI Salinity TA File EC02 RATE (meg/min) 111 04K 275.16 4000 36.248 2 .429 2.389 6 7 E-7 1111 AT II 14.6 2:431 2;389 2:4 E-7 11112K II 2075 II 2:436 2.392 4.0 E-7 11115K II 5080 II 2;443 2:396 2:03 E-6 1112AT II 14. 6 II 2:447 2.389 1 ;O E-7 1112QK II 2590 II 2:450 2:390 3:5 E-7 1112VK II 5475 II 2.456 2.393 2:05 E-6 Datasets # 1 0 (all 12 pts.) and #11 (first 6 pts.): Shell mass: 38.58 milligrams File Temp (K) PSI Salinity TA EC02 RATE (meg/min) 918ATM 275. 16 14.6 35.071 2.413 2.407 6.4 E-7. 918QK II 2475 II 2;429 2:415 1; 3 E-6 9185K II 4940 II 2:452 2:427 2:96 E-6 919A1M II 14.6 II 2:418 2:407 5:2 E-7 9192K II 2040 II 2:428 2:412 1.16 E-6 9194K II 3945 II 2:449 2:423 2:68 E-6 922A1M II 14.6 II 2.430 2.407 1 :45 E-6 922QK II 2550 II 2:458 2;42 1 1:47 E-6 9225K II 4910 II 2.486 2 .435 2:04 E-6 925A1M II 14.6 II 2;425 2;407 1:5 E-6 9251K II 980 II 2:453 2:421 1 :13 E-6 9252K II 2005 II 2:473 2; 431 1;15 E-6 Dataset /112: Shell mass -16.29 milligrams File Temp (K) PSI Salinity TA EC02 RATE (meg/min) DIS81 278.16 3100 34.681 2.513 2.425 8 6 E-7 DIS82 II 4960 II 2;531 2:434 1:43 E-6 DIS91 II 3150 34.606 2 .494 2:456 1:18 E-6 DIS92 II 4450 II 2:517 2;467 2:34 E-6 DIS93 II 1770 II 2.563 2:490 9.3 E-7 DIS1 01 II 3020 II 2:494 2:456 1 ;57 E-6 DIS102 II 4500 II 2 .527 2:472 3:53 E-6 DIS103 II 1770 II 2;594 2 :506 4;0 E-7 DIS111 II 2110 II 2:494 2 .456 1 .81 E-6 DIS112 II 4080 II 2;526 2 :472 3:34 E-6 DIS11 3 II 4970 II 2:59 1 2:504 4:3 E-6

PAGE 249

APPENDIX 2, continued 223 Dataset #13: Shell mass -16.29 milligrams File Temp (K) PSI Salinity TA l:C02 RATE (meg/min) DIS41 278.16 270 33.767 2.351 2.286 3.2 E-7 DIS48 II 300 33:766 2:356 2 :312 4;6 E-7 DIS49 II 1475 34:373 2.483 2:490 2:41 E-6 DIS50 II "730 34; 103 2:430 2:434 2:23 E-6 DIS51 II 520 33.696 2.380 2; 381 2:15 E-6 DIS51X II 2950 34:587 2:528 2:496 3:48 E-6 DIS52 II 2210 34.506 2 .508 2:502 4;26 E-6 DIS53 II 4320 34:656 2:528 2:460 4:66 E-6 DIS55 II 390 34.606 2:377 2:348 1 ;46 E-6 DIS401 II 15. 0 33:696 2 .380 2:381 8 9 E-7 DIS402 II 520 II 2;401 2:391 1.42 E-6 DIS47 II 710 34.049 2;402 2:383 1 : 91 E-6 Dataset 1114: Shell mass 43.38 milligrams File Temp (K) PSI Salinity TA rco2 RATE (meg/min) 23ATM 275.16 14. 6 36.268 2 .429 2.393 5.8 E-7 232K II 2100 II 2:439 2:398 1:51 E-6 235K II 4960 II 2 .465 2;411 3:59 E-6 24A1M II 14.6 II 2:423 2:393 5:1 E-7 24QK II 2595 II 2;432 2:397 1. 41 E-6 24VK II 5475 II 2:457 2 :409 3 .57 E-6 25ATM II 14.6 II 2 .438 2.393 2;5 E-7 25TK" II 3530 II 2;443 2:395 1:76 E-6 25VK II 5420 II 2:473 2 .410 3:1 E-6 26ATM II 14.6 II 2 :445 2:393 6:0 E-8 26FK II 1600 2 .446 2:394 1 .0 E-6 264K II 4170 II 2:462 2:402 2:72 E-6 Dataset #15: Shell mass (first 3) 19.2534 mg (all others) 18.257 mg RATE (meg/min) File Temp (K) PSI Salinity TA EC02 DIS161 278.16 15.0 34.539 2 .530 2 .475 2 6 E-7 DIS162 .. 2225 2:533 2:477 1:62 E-6 OIS163 II 4510 II 2 .569 2 .495 4.65 E-6 DIS181 4100 34.683 2:498 2;402 1:44 E-6 DIS182 5550 2:535 2:421 3 .52 E-6 DIS19 950 33.99 4 2:360 2;300 5:3 E-7 DIS201 II 15.0 34.305 2 .451 2:454 1 .53 E-6 DIS202 II 1530 II 2;480 2;469 2:75 E-6 DIS21 II 740 33.962 2:337 2. 291 9.8 E-7 DIS231 15. 0 34:024 2:355 2.333 9:0 E-7 DIS232 II 900 II 2 .377 2 .344 1;65 E-6 OIS24 II 1600 34.023 2;325 2:245 3:8 E-7

PAGE 250

APPENDIX 2, continued 224 Data 28: Shell mass: 33.51 File Temp (K) PSI Salinity TA tco2 RATE (meg/min) 106ATM 275.16 14.6 35.071 2.409 2.407 2.9 E-7 106QK II 2450 II 2;410 1 E-6 1 065K II 4940 2.433 2.419 2;69 E-6 107ATM 14.6 E-7 10715K 1500 II 1.13 E-6 1073K 3000 2;447 E-6 109ATM II 14.6 2 .419 1.16 E-6 1092K 2000 II 441 1 ;06 E-6 1095K II 4910 2.461 2;428 2;99 E-6 1011AT II 14.6 II 2;421 2.407 2;58 E-6 1011FK II 1525 2.467 2;430 1 :78 E-6 10114K 3920 II 2;499 2. 446 2;56 E-6 1012AT II 14.6 II 2;412 2;407 2;19 E-6 1012FK II 830 II 2;453 2;427 1 ;48 E-6 1012VK II 4375 2;479 2;440 3:10 E-6


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