University of Southern California Dissertations and Theses

Geochemical and sulfur isotope study of Red Sea geothermal systems

(USC Thesis Other)

Geochemical and sulfur isotope study of Red Sea geothermal systems

PDF

Download

Open document

Flip pages

Download a page range

Download transcript

Copy asset link

Request this asset

Transcript (if available)

Content GEOCHEMICAL AND SULFUR ISOTOPE STUDY OF RED SEA GEOTHERMAL SYSTEMS by Wayne Carlton Shanks, III A Dissertation Presented to the FACULTY OF THE GRADUATE SCHOOL UNIVERSITY OF SOUTHERN CALIFORNIA In Partial Fultillment of the Requirements for the Degree DOCTOR OF PHILOSOPHY (Geological Sciences) June 1976 UMI Number: DP28538 All rights reserved INFORMATION TO ALL USERS The quality of this reproduction is dependent upon the quality of the copy submitted. In the unlikely event that the author did not send a complete manuscript and there are missing pages, these will be noted. Also, if material had to be removed, a note will indicate the deletion. Dissertation Publishing UMI DP28538 Published by ProQuest LLC (2014). Copyright in the Dissertation held by the Author. Microform Edition © ProQuest LLC. All rights reserved. This work is protected against unauthorized copying under Title 17, United States Code ProQuest LLC. 789 East Eisenhower Parkway P.O. Box 1346 Ann Arbor, Ml 48106- 1346 U N IVER SITY O F S O U TH E R N C A L IF O R N IA T H E G R A D U A T E S C H O O L U N IV E R S IT Y P A R K ^ LO S A N G E L E S . C A L IF O R N IA 9 0 0 0 7 <y This dissertation, w ritten by WAYNE CARLTON SHANKS III under the direction of h.?r$... Dissertation C om mittee, and approved by a ll its members, has been presented to and accepted by The Graduate School, in p a rtia l fu lfillm e n t of requirements of the degree of D O C T O R O F P H I L O S O P H Y Dean D a t e . Q . Q l l d M ^ ^ ^ . . \ 3 . 3 S . DISSERTATION COMMITTEE ’ .airman G < 2 . S 52.8 CONTENTS Page LIST OF ILLUSTRATIONS............... v ABSTRACT.......................................... ix INTRODUCTION............... 1 General statement •••••••••••••• 1 Plate tectonics and ore deposits •••«•• 3 Objectives.......... # ................. 9 Acknowledgments ••••••••••••••• 11 PREVIOUS WORK . ............. 13 Historical note ................... 13 Bathymetry and hydrography of* hot brines # * l4 Chemistry of* hot brines ••••••••••• 27 Cold brine deeps •••••••••••••• 35 Interstitial brines « • ................. 37 Origin of hot brines •••••••••••• 37 Origin of cold brines •••••••••••• Metalliferous sediments ......... •••••• 51 Facies differentiation and origin «••••• 58 Sulfur isotopes ••••••••••••••• 63 ANALYTICAL PROCEDURES ............... 71 General statement .................... 71 Mineral separations ••••••••••••• 75 ii Page X-ray diffraction •.••••••••••••• 75 Sulfur isotope analysis 78 Chemical analysis ••••••• ........... . . 81 Radiocarbon determinations •••••••••• 82 RESULTS............................................... 84 Introduction •• •• •• ••• •• •• •• •• 84 Stratigraphic summary ••••••• • 87 Geochemistry of Atlantis IX Deep sediments • . 94 ¥ando, Albatross, and Shagara Deeps •••••• 107 Thetis, Suakin, and Gypsum Deeps • ......... 109 Other deeps and interstitial waters........... 113 Carbon-sulfur relationships •••••••••• 113 GEOCHEMICAL MODEL................................... 120 General statement •••••••* ............. 120 Conditions •• ••• • • • • • » • • • • • • • 121 D a t a ........................ 127 Calculations ••••• •• •• *•• •• •• • 133 Speciation ......... 136 Mineral solubilities .................. •••• l44 Sulfur isotopes and species distribution • . . 149 DISCUSSION AND CONCLUSIONS ........................ 154 Origin of Atlantis II sulfides .......... 154 Sulfur isotope ratios of sulfides ••••••• 137 Origin of Atlantis II sulfates.............. 170 Page Summary.......................................... 1?2 REFERENCES........................ 175 APPENDIX........................ 189 iv LIST OF ILLUSTRATIONS Figure Page 1. Location map of Red Sea deeps containing hydrothermal sediments . . . . ♦ ......... 5 2. Bathymetric map of the Atlantis II Deep area........................ . ............. 15 3* Temperature profiles in the Atlantis II Deep a r e a ................................. 18 4. Chlorinity profiles in the Atlantis II Deep a r e a ............. 20 5* Schematic diagram of convection and brine circulation in the Atlantis II Deep • • • 25 6. Major element composition of brines from the Atlantis II Deep a r e a ................ 30 7. Major element composition of cold brines • 33 8. Oxygen and deuterium isotope relation ships in some Red Sea waters............. 40 9. Oxygen and deuterium isotope relation ships of some thermal mineral waters . . . 42 10. Schematic diagram of facies relationships in Atlantis II Deep sediments • ••••• 53 11. Generalized stratigraphic sequence in the Atlantis II Deep........................... 59 12. Sulfur isotope ratio variations of some important natural compounds ♦ 65 13* Sulfur isotope distribution in Atlantis II Deep brines and sediments............. 67 v Figure Page 14. X-ray diffraction patterns of sulfide facies minerals • • • • • • • • • • • • • • • • • • 79 15. Detailed location maps for Suakin, Nereus, and Thetis deeps . ............. ............ 85 16. Three-dimensional projection of Atlantis II Deep stratigraphy • «.••••«•#••» 88 17. Lithostratigraphic units of cores 2 and 3 showing geochemical variations with depth • 91 18. Lithostratigraphic units of cores 6 and 8 showing geochemical variations with depth . 95 19# Lithostratigraphic units of cores 12 and 17 showing geochemical variations with depth . 97 20. Lithostratigraphic units of cores 19 and 20 showing geochemical variations with depth . 99 21. Lithostratigraphic units of cores 27 and 28 showing geochemical variations with depth . 101 22. Sulfur isotope fractionation between co existing sulfur—bearing minerals in the Atlantis II Deep •••••••••••••• 105 23* Variation of sulfur isotope ratio with organic carbon content of sediments • • • • 114 24. Variation of sulfur isotope ratio with carbonate carbon content of sediments . . . 116 25# Oxygen fugacity estimates for Atlantis II brine • • • • • • • • • • • • • • • • • • • 125 26. Dissociation constants for NaS0^“ ..... 128 27# CO^ activity coefficients utilized in geo chemical model ••• ••• •• •• •• ••• 131 28. Variation of hydrogen-containing species with temperature in the Atlantis II brine . 137 29. Variation of sulfur speciation with temperature in the Atlantis II brine • • • • 140 vi Figure Page 30, Variation of chloride complexes with temperature in the Atlantis IX brine , , , , Xb2. 31, Solubility of important hydrothermal minerals in the Atlantis II b r i n e .... 1^5 32, Sulfur isotope fractionation factors • • • • 150 33* Sulfur isotope ratios of sulfur species in the Red Sea b r i n e ...................... 159 3^# Variation of average sulfur isotope ratios in massive sulfide ore deposits with geologic time ............. 162 vii Table Page I. Major element composition of hot brines • • 28 XI, Minor element composition of hot brines • • 29 III. Chemical composition of cold brines • • • • 36 IV. Chemical composition of Leg 23 pore waters 38 V. Mineralogy of Atlantis II deposits • • • • 55 VI. Average composition of Atlantis II sediments • ••••••••••• 56 VII. Piston core data from Atlantis II Deep . . 72 VIII. Core data for areas outside the Atlantis II Deep...................................... 73 IX. Radiocarbon data ••••••••••••• 83 X. Geochemical data from cold brine deeps . • 110 XI. Dissociation constants of aqueous complexes 13^ XII. Degree of saturation of the hot brines . . 148 viii ! ABSTRACT i The Red Sea geothermal system provides a unique op portunity to study a submarine hydrothermal ore deposit I in the process of formation. Within the Atlantis XX i ideposits, massive sulfide layers are of particular J ' importance because of compositional and isotopic simi- !larities to ancient deposits. ! Re-examination of stratigraphic data on the Atlantis 1II deposits, combined with radiocarbon age dates, indicate that basin-wide correlation of facies is not possible. A complex and discontinuous history of brine activity is revealed wherein sulfides are differentially precipitated near the brine vent. Sulfur isotope ratios of sulfide minerals within the Atlantis II Deep vary from +15 to -45 °/oo. Values rang ing from -20 to -45 °/oo are ascribed to bacteriogenic sulfate reduction when brine activity is minimal. Base- Qlj. O / metal sulfides have S S ratios of 0.0 to +15 /oo and j are directly related to hot brine activity. Sulfide is i • a result of sulfate reduction in the brine and sulfate— i |sulfide fractionation is controlled by kinetic isotope effects. Similar hydrothermal sulfides are found outside r — \ iof the Atlantis IX Deep area in the Thetis and Shagara |Deeps. Geochemical modeling of brine evolution during sub- isurface circulation indicates that NaSo^~ and MgSO^° are the major sulfur-bearing species, thus maintaining a : uniformly low concentration of l+2‘ 5 (/^l ppm) in the brine. i Sulfide minerals are precipitated during cooling due to :release of free metal ions from chloride complexes. Estimations of brine flux and sedimentation rates indicate agreement of this model with observed sulfide mineral i concentrations in the metalliferous sediment. x j INTRODUCTION General Statement The discovery and subsequent investigations of hot brines and highly metalliferous sediments in the axial rift zone of the Red Sea (Degens and Ross, 1969) have revealed a submarine hydrothermal ore deposit which is still actively forming. The metalliferous sediments in the Atlantis II Deep area in the central Red Sea may prove 1 to be an economic mineral deposit, if political and extractive problems are ever overcome. Even more exciting from a scientific point of view is the opportunity to study an ore deposit in the process of formation. An obvious limitation on the study of ancient metallic deposits is the almost complete lack of direct information on ore-forming fluids. Virtually everything !that is known about ore-fluids is inferred from mineral equilibria (Holland, 1965)* theoretical geochemical considerations (Helgeson, 1964), or compositional data jderived from fluid inclusion studies (Roedder, 1967)* i The first two methods have provided important general information and the latter studies have delineated some 1 compositional and physical restraints. An obvious, and perhaps serious, limitation of fluid inclusion studies is the requirement that the fluid be a representative sample of the ore fluid which is physically trapped in the rock and has not reacted extensively after isolation. Ana lytical procedures are also a problem, especially in the determination of trace amounts of the important ore- forming metals. Thus, the Red Sea hot brine pools provide an important opportunity to study in considerable detail an ore-fluid which presently is active. For the first time parameters such as total dissolved solids, trace metals, isotopic ratios, temperature, and discharge rates can actually be measured. Study of the Red Sea geothermal system also may provide evidence germane to the old argument of syngenetic versus epigenetic formation of massive sulfide deposits. Stratigraphic and textural studies of the metalliferous sediments, coupled with studies of particulate matter precipitated in the brine deep, leave no doubt as to the syngenetic nature of this deposit. What remains somewhat obscured is the exact relationships of this deposit to grossly similar ancient deposits. Certainly detailed mineralogical, chemical, and isotopic studies, all essen tial objectives of this study, are most important in 2 j relating the metalliferous muds to their lithified counter- iparts, Plate Tectonics and Ore Deposits The formation of metalliferous sediments at active oceanic ridges is well documented, Bostrom jet al. (1969) and Sayles and Bischoff (1973) have described iron- oxy-hydroxide and iron-silicate-rich sediments from all of the active ridge systems. In addition, Bostrom at al. :(1972)f von der Borch and Rex (1970)> von der Borch at al. (1970), and Cronan at al. (1972) have studied similar iron-oxide-enriched sediments from Deep Sea Drilling Project cores in the Atlantic and Pacific oceans. These metalliferous sediments overlie basal basalts and presum ably were formed at ancestral ridges and later buried. Numerous authors, including Bischoff and Dickson (1975)» have suggested that these sediments are deposited from sea water which is heated and circulated at depth through fractured basalts at the ridge crests. The important features of this model are the high heat flow near the ridge crest which acts as a driving force for sea water convection and basalt-sea water interaction to provide a source of heavy metals. The Red Sea geothermal systems differ from open ocean ridge systems in several important aspects. The Red Sea provides an excellent example of rifting and break-up of :continental crust to form an embryonic ocean basin. Mafic i ;intrusive rocks have been emplaced along the length of the axial rift zone and seismic refraction studies (Drake and Girdler, 1964, and Phillips ejfc ad*, 1969) suggest that continental crust may underlie a large portion of the Red Sea flanking the axial rift zone* Above these continental rocks is a thick sequence (3 to 6 km) of evaporites typical of early continental break-up, restricted ocean basins, and arid conditions. Leg 23 of the Deep Sea Drill ing Project has sampled these evaporites at sites 225, 227, and 228. They consist of remarkably clean sequences of anhydrite and halite which are continuous to the edge of the active rift zone. Bignell (1975) has suggested that the formation of depressions of suitable geometry for deposition of metal liferous sediments is related to transform faults which cut across the Red Sea* According to his model, transform faults provide conduits for metal-bearing fluids to circulate through impermeable evaporite sequences. Where transform faults intersect the median valley a metal liferous deposit or brine pool may form (Fig. l). The geologic situation in the Red Sea provides a number of distinct advantages to the generation of metal liferous deposits, especially base metal sulfides. As along open ocean rift systems, the active intrusion of mafic volcanics at the rift axis provides a heat source Figure 1 Location map of Red Sea deeps contain ing hydrothermal sediments (after Backer and Schoell, 1972)# RED SEA EGYPT HOT BRINE AREA VALDIVIA SUDAN SUAKIN DEEP KEBRIT DEEP GYPSUM DEEP VEMA DEEP NEREUS DEEP THETIS DEEP HADARBA-HATIBA DEEP ALBATROSS DEEP SHAGARA DEEP SAUDI ARABIA YEMEN ETHIOPIA 6. | for potential fluid circulation* In tine Red Sea rift, j j however, bathymetry is quite restricted and closed basins | I are readily available. The presence of evaporites and ; abundant sediments in close proximity to the rift appears i tto be especially important. 1 i Circulating waters may interact with evaporite t minerals (especially halite) to form dense brines which :tend to pool in the deeps and result in localized deposi- ! tion of metals. In terms of metal transport, organic- ‘rich shales in sedimentary sequences provide a potential metal source and complexation by chloride in concentrated brines greatly increases solubilities of ore metals. [ Tectonic situations similar to the present Red Sea would be expected to have existed at various times in the ■ geologic past when break-up of continents was occurring. An excellent example is the formation of the proto— , Atlantic Ocean. Blissenbach and Fellerer (1973) have discussed the possible occurrence of submarine hydro- ' thermal sulfides in the Atlantic. They conclude that the embryonic North Atlantic in Jurassic time probably | exhibited ideal prerequisites for generating metal I deposits: restricted marine basin, flanking evaporite jsequences (Schneider and Johnson, 1970)* and active rift- i 1 ing. To date, however, only two occurrences of mineral- I ization have been reported, both from Deep Sea Drilling ! Project sites in the northwest Atlantic. Site 105 of Leg I 11 sampled about 50 hi of* late Cretaceous ferromangoan sediments which, contain locally abundant Cu, Fe, and Zn sulfides. These sediments are similar to the Red Sea !deposits but occur in an off-ridge type environment, over- lying 350 m of Jurassic-Cretaceous sediments. Nonethe less, this discovery confirms the potential universality of submarine hydrothermal activity associated with divergent plate boundaries. There are numerous examples in the literature of ophiolite—associated massive sulfides which may have formed by processes similar to those operative in the Red Sea. A complete discussion of these deposits is beyond the scope of this brief introduction, however, a few out standing examples can be mentioned. Hutchinson and Searle (l97l) and Constantinou and Govett (1973) have described the massive sulfide deposits of Cyprus in some detail. These lensoid, pyrite— chalcopyrite-sphalerite bodies overlie pillow lavas and other volcanic rocks which have been interpreted as hav ing formed at a Tethyan oceanic ridge (Moores and Vine, 197l). In stratigraphy, general geometry, and gross composition, these deposits resemble the Atlantis XX Deep sediments. The massive sulfides at Cyprus are underlaid by stockwork "feeder” veins of sulfide and grade upward into oxidized and silicic zones which contain only dis seminated sulfides. The close similarity with the Atlantis 8 |IX deposits suggests similar depositional environments* | Another important example of* similar massive sulfides is from the Newfoundland ophiolite belt* Upadhyay and i ; Strong (1973) have described occurrences at Betts Cove |and Duke and Hutchinson (197^) have studied those near !York Harbour* In general they are Cu-Zn ores similar in many respects to the Red Sea deposits* Obviously, detailed unraveling of processes active in the Red Sea is 'important to the interpretation and location of such 1 similar ancient deposits* In addition to ophiolite-related massive sulfides, t Anderson (1969) and Hutchinson (1973) have synthesized information on similar deposits in other types of volcanic rocks. In fact, numerous authors have proposed classifi cation of massive sulfide ore-types according to plate 1 / tectonic environment (Sawkins, 1972; Sillitoe, 1973; Solomon, 197^1 Hutchinson, 1973)* It has been suggested that a number of massive sulfide deposits in andesitic or rhyolitic volcanoes may have formed by re-circulated, | heated sea water. Thus the geochemical processes occurring i i in the Red Sea may have widespread analogies among many ! different types of massive sulfide deposits. ! I Objectives 1 ! Although considerable geochemical work has been done already on the Red Sea geothermal system (Degens and Ross, _______________________________________________________________________ L_ 1969)> there are some persistent and important problems still unresolved. Recent stratigraphic studies (Hackett and Bischoff, 1973; Backer and Richter, 1973) have revealed ;that the sulfide-rich zones in the Atlantis II Deep are much more extensive than originally sampled. This obser- 1vation has strengthened the analogy to ancient deposits and provides impetus to further study the genesis of the 1base-metal sulfides. Earlier studies using sulfur !isotope geochemistry suffered from inadequate sampling of :the sulfide facies and have proved inconclusive. One major purpose of this study was to delineate the source I of the sulfur and the mode of precipitation of sulfide i j facies in the Atlantis II Deep. 1 The methods chosen were: detailed re-examination of 1 stratigraphy and geochronology of the deposit, systematic I analysis of sulfur isotope ratios in carefully separated phases from the sulfide facies, and analysis of sulfur, organic carbon, and carbonate carbon in relation to sulfur isotope ratio. These methods have proved to be powerful tools in determining the genesis of sulfur-bearing minerals in recent marine sediments (Goldhaber and Kaplan, 197*0 and in ancient ore deposits (Rye and Ohmoto, 197*0* A second unsolved problem in the Red Sea system is the evolution of the Atlantis II brine. There has been much controversy in the literature over the exact source of the waters (Craig, 19^9; Ross, 1972; Manheim, 197*0 and 10 little evaluation of the brine’s potential to transport ore# As more data have become available on the hydrog raphy of the system (Schoell, 1975) it has now become possible to quantitatively evaluate brine evolution. The method chosen here is an aqueous complexing model which |is used to calculate distribution of dissolved species and to evaluate solubility with respect to important ore :minerals over a range of conditions. I • • I The discovery (Backer and Schoell, 1972) of many cold ,brine deeps in the axial rift zone leads to the third I emphasis of this study. Sediment samples were obtained 'from all the deeps known to contain hydrothermal sediments 'and mineralogical, geochemical, and isotopic studies were I carried out to determine possible genetic relationships I or similarities to the Atlantis IX deposit. i Acknowledgment s It is a pleasure to express my sincere gratitude to 1 James L. Bischoff, my committee chairman, who has made ,innumerable contributions to the successful completion of i this project. He has provided guidance, encouragement, A and many useful ideas. The quality of the final manuscript !owes much to his careful reading and rereading. t \ I also wish to thank Dr. I. R. Kaplan, who kindly !provided complete access to the laboratory facilities at r ' 1 1 - - - - - - - - - - - - - - - - - - - -- - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - - i i the University of California at Los Angeles and to many |years of sulfur isotope expertise. Special thanks are due to Dave Winter at U,C,L,A, for considerable laboratory assistance, | Useful ideas and critical comments on the manuscript ■ were shared by Richard Ku, K, Y, Chen, Robert Sweeney, Tom O'Neil, Tom Nardin, Kevin Knauss, Martin Schoell, and imany others, Kerry MacLennan provided expert laboratory i ! and "field” assistance. Special thanks are owed Tom , O'Neil for computer assistance. My wife Linda typed several early versions of the t manuscript, served as chief managing editor, and provided I ; moral support throughout, as did Richard Stone and Donn Gorsline. I m m I Finally, X wish to acknowledge Dr. Harold von Backer i i of Preussag AG for kindly providing many sediment samples. This study was supported by NSF Grant GA 23^95* 12 I I ! PREVIOUS WORK i ! Historical Note Warm, highly saline bottom waters in the central Red Sea were recognized as early as 1948 by the Swedish Albatross Expedition (Bruneau et al,. 1953)* Another !important early discovery of4 anomalous water was made by the Woods Hole ship, Atlantis (Neumann and Densmore, 1959)* However, the significance of these reports was not recognized until the hot brine area was investigated by a number of ships passing through the Red Sea while participating in the International Indian Ocean Expedition i (1963-1965). In 1963 the Atlantis II, occupying routine hydro- graphic stations along the length of the Red Sea, re discovered the anomalous waters (Miller, 1964, 1969)* Subsequent conformatory observations were made by crews ,of the Discovery (Swallow, 1969)9 the Meteor (Dietrich and 1 i Krouse, 1969)9 the 0ceanographer (Ostapoff, 1969)9 and the Akademician Vavilov (Shishkina and Bogoyavlenskiy, 1970), ; In addition, comprehensive geological, geophysical, and 1 geochemical investigations were carried out in 1966 by the 1 , Chain cruise 6l* The results of this cruise were pub- : 13 lished as a symposium volume (Degens and Boss, 1969) which, remains the most complete source of* information on the Red Sea geothermal system. I ! More recently, interest in the base-metal sulfides of the Atlantis XI Deep as a potential submarine ore !deposit has stimulated industrial participation. Detailed 1 ,sampling and exploration of the Atlantis II Deep has been i achieved by the Wando River in 19^9 and the Valdivia in i ; 1971 and 1972. i i Bathymetry and Hydrography of Hot Brines The geothermal brines of the Red Sea occur in several * discrete or partially connected pools within the axial irift zone approximately due west of Jidda (Fig. l). The 1 ; largest of these pools, the Atlantis II Deep, measures 1 3 '5 km by 14 km and contains approximately 5 km of super saline brines (Schoell, 1975)* To the immediate south (Fig. 2), and apparently connected with the Atlantis II Deep, are the Chain Deeps. Designated Chain Deeps A, B and C by Backer and Schoell (1972), these basins are ap- |proximately 1,5 hm in total free brine-surface area. The Discovery Deep to the southwest of the Atlantis II Deep i 2 has a brine-surface area of about 11.5 hm . i In addition to the well known brine pools just 1 mentioned, the brines of the Wando Terrace and the Albatross Deep are included in this section. It can be 14 Figure 2. Bathymetric map of* the Atlantis IX Deep area (contours in meters). Stipled areas indicate approximate extent of* brine pools. Sample locations are from Wando River and Valdivia cruises (after Backer and Richter, 1973). ATLANTIS WANDO AS/N^i 2r20 VALDIVIA QE;E/* 4 CHAIN DISCOVERY DEEP 2 1 1 5 2r 10 SHAGARA 2 1° 2 5 ' M 37°55' 38°oo' 3 8 ° 0 5 / 3 8 ° 1 0 ' j ; inferred from bathymetric proximity (Wando Basin) and |brine composition (Albatross Deep, see below) that these i | brines are genetically linked to the Atlantis II geo- i j thermal system. j Figures 3 and 4 (Schoell, 1975) summarize temperature- !pressure and chlorinity-pressure profiles in these deeps obtained during the Valdivia cruises. Pressure is plotted instead of depth because of difficulties in estimating jsound velocity in the various brine layers. The Atlantis i II Deep is divided into two major hydrographic units with i several minor and transitional brine layers. The lower |Atlantis II brine has a maximum temperature of 59*8° C (1972), a chlorinity of 156.5 °/oo, and occurs below the 2040 m contour. The upper or transitional water mass has a temperature of 50° C, chlorinity of 82.5 °/°o, and occurs roughly below 1990 ni. It has been established that this brine layer is a result of mixing of the lower brine mass with normal Red Sea Deep Water (Craig, 1969? Turner, 1969). An important result of the Valdivia exploration is the discovery of high chlorinity (155*5 °/oo), high temperature (51*6 C) brines in the Chain Deeps. The |maximum previous temperature detected was 34° 0 (Ross, '1969)» undoubtedly due to the difficulties in placing a hydrocast in such a small basin. Discovery Deep (Figures 3 and 4) has a maximum 17 Figure 3« Temperature profiles in the Atlantis II Deep area (Schoell, 1975)* * 1950 2000 2050 2100 2150 2200 2250 - 2300 20 30 AO 50 60 1----------1 --------1 --------- f l --------1 TEMPERATURE (#C) 22.0 RED SEA DEEP WATER ,50 .6 (49,8) ,50,9 ,51.4 (50,2) MAX. 60,09 (59,76) WANDO-BASIN 29,4 *51 . 6 CHAIN-A A L8ATR0SS-DEEP 24,4 CHAIN- B 46.10 S'//- BASIN 59,68 (59,20) ATLANTIS I DEEP DISCOVERY DEEP 44,80 NORTH-BASIN 05 < CD Q U l LU 05 19 Figure k Chlorinity profiles in the Atlantis XX Deep area (Schoell, 1975)# 1950 2000 2050 2100 2150 2200 2250 j 150 154 155 155 157 |V l— I — I — 1 CHLORINITY Cl C7oo] ATLANTIS 11-DEEP 61 7 %• 1 838%. J UPPER B R IN E (50*C ) ALBATROSS-DEEP ATLANTIS 11-DEEP ' s ^ ' 1 5 6 ' 5 %• LOWER BRINE (5 9 ,6 *0 • CHAIN - C / ° CHAIN- B 7 5 5 ,5 % . * n cr < co o DISCOVERY-DEEP 155,0 V- A 755,4 % 21 j temperature of4 44.8° C, and a chlorinity of* 155*0 °/oo. j Pugh (1969) asserted that the Discovery Brine is a result |of overflow from the Atlantis II Deep because temperature ;profiles indicated cooling from the bottom as well as 1 I ;from the top. This basal temperature reversal was not 1 duplicated by the Valdivia exploration even though several ;detailed temperature profiles were obtained (data compiled by Schoell, 1974a and b). A striking feature (Pigs. 3 and 4) is the concurrence of depths (in this case pressure) to the tops of the ihottest brines in the Atlantis II, Chain, Discovery and Albatross deeps. It is possible that this depth leveling !is a result of subsurface connection of all of these deeps, iIf this is the case, however, it is difficult to explain 1 the temperature differences and the occurrence of Wando I Basin brine at a markedly shallower depth (Fig. 2). In addition, Brewer et al. (l97l) reported hot saline brines in a narrow elongate depression to the northeast of the Atlantis II Deep (Fig. 2). This basin is separated I from the Atlantis II Deep by a sill depth of less than 1 !1900 m and maximum depth in the basin is less than 1980 m. Nonetheless, this basin contains a lower brine (56.7° C, j salinity 318 °/oo) and an upper stratified brine (49.9° C, salinity 209 °/oo. i Continued observation of the temperature of the Atlantis II brine since 1966 (Brewer ejb al., 19715 Ross, ! 22 il972; Schoell and Hartmann, 1973; and Schoell, 1975) has j indicated an increase in temperature in the lower brine 1 o j mass of* 0.75 C per year. Since the change in brine temperature is accurately known (i 0.05° C) the temperature of* new brine input can easily be calculated if* the volume of* new brine can be discerned. . ! Brewer jet al. (l97l) and Ross (1972) calculated brine input temperatures of about 110° C based on a supposed 6 m I rise in the depth of the lower brine layer between the 1966 and the 1971 Chain cruises. These estimates are subject to considerable doubt, however, because they are based on depths determined by the length of hydrowire paid out and an estimate of wire angle. Brewer e_t al. (1969) 1 stated that the 1966 Chain deeps are accurate only within ■ - 5 m and, therefore, no reasonable temperature estimates can be calculated. Schoell (1975) provides a more probable estimate of the temperature of brine input by comparing depth data from the 1964 Meteor cruise and the 1971-72 Valdivia |cruises. As he points out, both vessels were equipped I with the same high-precision ELAC narrow-beam (l.4°) echo sounding systems. Direct comparison of echo sounder I |recordings indicated the exact same depth to the lower brine layer in 1964 and 1971* Schoell (1975) estimates the accuracy of the ELAC systems to be within 2 m and calculates a brine input temperature of 210° C accordingly. 23 It should be emphasized that this is a conservative j estimate of input temperature for two reasons: (l) it is j based on a 2—m increase in the brine pool between 19^5 and ,1971$ when it actually may have increased by less; (2) the I inflowing' brine should have experienced some heat loss to ; the surrounding sediments# t I If the volume increase is less than 2 m, then the jinput temperature must be proportionally higher. A )reasonable upper limit to brine input temperature can be f f I estimated from silica solubility data. Kennedy (1950) has r tshown that there is a significant reversal in silica t |solubility at about 350 C, for pressure less than 700 bars. Decreases in silica solubility in this region would ,result in quartz precipitation which would rapidly seal 1 off supply conduits. i • The detailed temperature profiles obtained by the iValdivia cruises have allowed Schoell and Hartmann (1973) and Schoell (1975) to compile considerable detail about the present source and circulation of the lower Atlantis II brine. Representative temperature profiles from I various sub-basins are given in Figure 3* As summarized in Figure 5 (Schoell, Backer, and Baumann, 1975)» the !present brine vent is in a small area of the southwest 1 basin, as previously suggested by Bischoff (1969) on the !basis of mineralogy. From its source the brine convects 1upward to the lowermost brine interface and then spreads 2k Figure 5. Schematic diagram of* convection and brine circulation in the Atlantis IX Deep (Schoell, Backer, and Baumann, 1975)* WATERDEPTH IDBAR1 £ ^ £ ^ f3 i l i laterally into the west basin, east basin, and north basin* Chemistry of4 Hot Brines The chemical composition (Tables X and XX) of the Atlantis II and Discovery Deep brines have been determined by a number of investigators in recent years, notably Brewer ej; al. (1965)* Brewer and Spencer (1969)$ and Man- heim (1974). Most elements are enriched in the brines, especially Na, Ca, K, and Cl, but notably depleted with respect to normal sea water are Mg, I, SO^, NO^, and P. OT particular importance are Pe, Mn, Pb, Zn, and Cu which are very greatly enriched over normal sea water and are obviously very important in the ore-forming processes. As demonstrated by Craig (1969) and Brewer and Spencer (19^9)$ major element composition provides much information about the origin of the various brine- related water masses. A plot of major elements versus chlorinity for the various waters (Fig. 6) shows that the transitional brine layers plot on linear mixing lines between Red Sea deep water and the Atlantis II 60° C brine, indicating conservative behavior. In addition to the transitional waters (Atlantis II 44°, Chain 34°), waters from the Shagara and Albatross deeps are also plotted. The Albatross Deep water analysis is courtesy of Preussag AG (H. Backer, written communication). The 27 ! Table I. Major element composition of* hot brines I (g/Kg) Element Red Sea Deep Water Atlantis IX Deep Discovery Deep Cl 22.5 156.3 157.3 155.3 Br 0 .0 7 6 0.128 0 . 119- 0 .1 2 3 so4 3.14 C O • 0 0.7 5 - 0 . 7 0 Total CO^ 0.1 4 0 . 0 7 5- 0 .1 3 2 0 .0 3 2 Na 1 2 .5 9 2 .6 9 2 . 8- 9 3 . 0 K 0 .4 5 1.87 2 .1 5 Ca 0 .4 7 5.15 4 . 7- 5 .1 Mg 1.49 0.764 0 .81 Total Dis solved Solids 4o,6 257.3 256.4 i i i 28 I Table IX# Minor element composition of hot brines (mg/kg) Red Sea Atlantis XI Atlantis II Bottom Deep Discovery Deep Element Water (560 C ) Deep (44° c) i As 0.1a 0.1a ; B 5.3 £ O • O H 7.5 Ba o.9b 0#3b Br 76.0 128.0C 119-123° jc° 0.16° 0.13° 0.0008° | Cu 0.26° 0.08° 0.017° F 1.6 O# 02e 0.051° Fe 0 O « H 00 0.27° 0.2° Li 0.21 4.i-in7d 4.0-4.ld Mn 82.0° 54.6° 82.0° Mo 0.01 S z j 0 0 1 0#001b o.ooib o.ooib NO^-N 0.11b o.onb o#ooib ; Ni o#34e 0.0012° : Pb 0.63° 0.009° 0.0088° ft 1 " ■ S O ft 0.02b O.02b o.oib Se 0.15a 0.05a SiO^-Si 0.3^b 12.3b 15.7b !Sr 9.2 48.0° 46. oc 27.0° u 3.3 0.it8f , Zn 5A° 0.77° 0.15° 'a Kaplan et. al# (1969). b Miller et al. (l966)# c Brewer and Spencer (1969) id Brooks et_ al# (1969)* e Brewer et al. (1965)* f Ku (1969T. 29 Figure 6. Major element composition of* brines from the Atlantis XX Deep area show ing mixing relationships between brines and normal Red Sea water. Data from Brewer ert al. (1965)9 Brewer and Spencer (1969)9 Craig (1969)9 and Backer (19759 personal communication). Chemistry of Shagara brine determined (this study) on pore waters. 30 □ 0 0 <> DISCOVERY DEEP ATLANTIS II 56° (1966) ALB A T R O S S ATLANTIS II (19661 CHLORIDE (g/ kg) 31 j Shagara brine composition was analyzed on two samples of* i pore water, correcting to a chlorinity of 113 /oo as |reported by Backer and Schoell (1972). It appears from Figure 6 that the Albatross brine is !genetically related to the present-day hot brine system while the Shagara brine is not. Although the present iShagara brine plots fairly closely to the mixing lines for ^ " t " " t " Na , K , and S0^=, the Mg concentration is much higher ithan even normal Red Sea deep water. This is suggestive ;of cold evaporite interaction and, in fact, this water ! composition is more similar to the cold brines and pore i waters discussed in the next section. I As mentioned previously, the mode of connection of t lthe Atlantis XX Deep with nearby brine deeps is somewhat ♦ t !uncertain. It is likely, however, that one or more over- l i |flow events are responsible. Figure 2 shows that an over flow would require an increase in the brine pool to a depth near 1900 m. Once accomplished, a brine overflow 'would result in density flow along topography past the Chain deeps, southward in a narrow trough to the Albatross ;Deep, and possibly spilling over into the Shagara Deep. ;If a recent overflow did reach Shagara Deep, it is dif- t ficult to explain the persistence of brine composition m 1 the Albatross Deep but not in the Shagara. 32 Figure 7« Major element composition of cold brines. Trend lines indicate chemical variations for isothermal evaporation of sea water* Data from Backer and Schoell (1972). C O N C E N T R A T I O N 12 GYPSUM P R E C IP IT A T E S 11 10 9 8 3 2 KEBRIT 1 20 10 G O 8 0 10 0 120 MO 1 6 0 CHLORIDE ( g/ kg) 3k i Gold Brine Deeps f i Backer and Schoell (l972) have reported the existence i of* numerous additional deeps in the axial rift area of the ;Red Sea (Fig. l). Many of these deeps contain cold, high- i salinity brines with appreciable concentrations of metal liferous sediments. None, however, contain metal concen- i Itrations comparable in volume and grade to the Atlantis XX deposits. In addition, most of the high-salinity brines |have distinctly different chemical compositions from the 2 + - hot brine area, lacking the characteristic Mg and S0^~ depletion (Table III)• ' Concentrations of major dissolved constituents in the cold brine deeps are plotted against chloride concentration ■in Figure 7« For reference, Red Sea deep water is |included. Sloping lines represent simple evaporative concentration of Red Sea deep water. For this range of chlorinities, the only solid phase considered is gypsum precipitation at 118 °/oo Cl. Earlier CaCO^ precipitation is considered quantitatively unimportant in controlling 2+ 'Ca concentration. Perusal of this diagram xndxcates jthat the cold brines have not formed by monotonic |evaporation of Red Sea waters and are, therefore, not i !simply residual evaporated waters from the last glacial i low-stand of sea level. i 35 i I Table IXI. Chemical composition of cold brines (g/Kg) 0 0 Na K Mg Ca so4 Cl Suakin Zh.j 4? 1.1 1.4 2.1 3.2 85.9 Port Sudan 36.2 78 1.5 1.5 1.1 4.0 125 Erba 27.9 50 1.0 1.5 1.0 4.1 86.5 Shagara* - 62 2.0 2.5 4.8 1.2 113 Albatross 24.4 86 2.0 1.0 4.0 1.1 143 Wando 29.3 - - - - - 73.5 Valdivia 29.7 95 1.9 1.9 0.8 6.8 140 Thetis 22.6 11.3 0.46 1.3 O • 00 00 3.^ 23 Nereus 30.2 - - 1.5 7.9 0.9 130 Kebrit 23.3 95 1.5 2.4 7.1 2.2 153 * Analyzed by ¥. E. Seyfried. i i I 36 Interstitial Brines I I j During Leg 23 of the Deep Sea Drilling Project in the i Red Sea, Miocene evaporites were penetrated in sites 2 25» 1227, and 228. Stoffers and Kuhn (197*0 Lave described evaporite mineralogy; mainly halite and anhydrite with j subordinate late—stage minerals such as tachyhydrite (Ca Mg2 Cl6 . 12 H20), polyhalite (K2 Mg Ca2 (SO^ . 6 H20), !and carnallite (K Mg C1Q . 6 HQ0). They concluded from j J ;textural and geochemical evidence that the deposits sampled (actually only the uppermost sections of evaporite sequences up to 3 km thick) represent the waning stages of i evaporite deposition in a highly dessicated basin, i Manheim (197*0 analyzed major elements in interstitial waters from the evaporites and younger overlying sediments. !The pore waters range in chlorinity from 22.5 °/00 (normal Red Sea water) to 155 °/00 (Table IV). These high chlorinity waters are strikingly similar in composition to the previously discussed cold brine pools. As demonstrated by Manheim, compositional consideration of added salts suggests that these interstitial brines are a result of interaction of connate waters with evaporite minerals. ■ Origin of Hot Brines The most compelling arguments concerning hot brine origin in the Red Sea have been put forth by Craig (1966, i 37 Table IV. Chemical composition of Leg 23 pore waters (g/Kg) _o „ T C Na K Ca Mg S4 HC03 Cl pH 225-10 23 12.6 0.46 0.49 1.47 3*62 0.12 22.5 7.7 225-45 26 14.3 0.44 0.77 1.36 3.37 0.10 25.3 7.9 225-70(64) 28.5 14.4 0.37 0.82 1.26 3.25 o.i4 25.2 7.6 225-130 33 24.0 0.30 1.10 1.01 3.08 0.12 39.0 7.2 225-220(212) 4l 91.7 1.25 0.67 2.14 4.22 0.06 148.3 -(7.0) 196 m 228-6 22.5 16.3 0.45 0.71 1.50 3.46 0.10 28.5 228-45(36) 31 21.9 0.38 0.94 1.45 3.36 0.13 37.7 7.3 228-85 37 34.1 0.35 1.44 1.46 3.26 0.13 57.2 - 7.0 228-190 a/ 80(60) 50.4 0.32 2.65 1.24 2.43 0.05 81.9 6.3 228-306 120(90) 66.4 0.52 1.20 1.46 2.75 0.08 107.5 7.0 0 00 18 1969). His c£*O and < 6 D determinations on the Hot brines, normal Red Sea waters, and other local waters are repro duced in Figure 8. Additional new data on the Atlantis XX Deep transitional brine, Wando, and Albatross brines have recently been collected by Faber (1975) and are also included. The Red Sea hot brines plot (Fig. 8) directly on the locus of* points delineated by S D-cJo^ relationships in normal Red Sea waters. This indicates that either the brine represents isotopically unaltered Red Sea water or that a remarkably fortuitous situation has occurred. Two alternative processes could derive the isotopic composi tion of the hot brines from meteoric waters. One such process is evaporation of meteoric water and the other is the oxygen isotope shift due to interaction with rocks (Fig. 9) commonly observed in continental geothermal waters. Obviously, either of these processes would require a coincidence to plot on the Red Sea waters line. An additional possible process has been suggested by White (197*0 whereby Saudi Arabian spring waters (which plot on the meteoric water line at <f D=-25 °/oo) could mix from 10—40 percent with local Red Sea waters. This mechanism, however, still concludes that the majority of the water is unaltered Red Sea water. Craig (1966, 1969)9 assuming the hot brines to be isotopically unaltered Red Sea water, uses D-salinity 39 Figure 8. Oxygen and deuterium isotope relationships in some Red Sea waters. Data from Craig (1 9 6 9 )* Faber and Schoell (l975)» Lawrence S 197*0 9 and Friedman and Hardcastle 197*0. +20 +10 8D (o/oo) 0 -10 -20 / / fc] 225- 45 PO 22 5- 70 228- 45 n r ^ fe o B 2 2 5- : V A L D IV IA BRIfE A R E A (in clu de s Aruvms II, D iscovery Chain mo A lbatross beeps) 228-3060 O D S D P P O R E W A T E R S O C O L D BRINE D E E P S -3 -2 -1 0 + 1 + 2 +3 SO18 ( c / o o ) 41 Figure 9* Oxygen and deuterium isotope relationships of some thermal mineral waters (White, 1968). p e r mil 6 0 '\ per mil - 1 5 - 1 0 - 5 0 + 5 + 1 0 + 1 5 ,- - - - - - 1 - - - - - - - - - - - - - - - - - - - - - - - - r - - - - - - - - - - - - - - - - - - - - - - - - , - - - - - - - - - - - - - - - - - - - - - - - - r - - - - - - - - - - - - - - - - - - - - - - - , - - - - - - - - - - - - - - - - - - - - - - - - - 5 -- ^ - R e c o r d e d range. 6 0 of fluid inclusio ns, ^ M s u s s i p p i Veil ey (Hell ond Friedman, I96J) V.'ciro’ -ei, _ Ne w 7eo!ond * O 00 - 10- - 1 5 Ocean - ( S i v ' . O W ) y \ x- Br.ne, M ic h ig a n copper . G , mir.e ( C ro tj e n ctyv s ' Red Seo ge othermal /. Nor mol Red Seo w c t e r s (C ra ig I 0 C 6 ) mir.e I L r o t j e n c:yv s . of White's sample} / + + + + Y- r~ ^ ♦ ++T / + -• Lcrdereiio, Italy + Illinois o ilfield w a t e r s (C la y t o n e n d e t h e r s , 19 G£>) ’ T h e G eysers”, Coiitornio a >> / H e h lo , Ic e la n d / r t f ' ^'v5>V & / .A / AX _______________________ / Sollcr, S e o g e o tr .e r m o l c r e o . C o M . M o g m c tic w c t e r *? 6 D cod calculated P . ^ y j 4 6 O ’, ® P r o v i d e n c i o U (R y 6 , 1 9 6 6 ) * M organ Springs, Mt. Lossen, Colif S ie o m b o o t Sprin gs, N e v o d o A. / U p p e r Eosin — Yellowstone Perk (C rc ig ’s analysis of W h ite s sample) Meteoric thermal X - ------------ • Trerd lims of s'jdied hydro- t n e r m o i s y s t e m s , showing I im il j of measured d il I - * rerces Modified irom Craig, 1967 43 r --------------------------------------------------------- I 18 |and 6 0 -salinity plots to infer an initial salinity of I I |38*2 /oo for the derived water# Because such waters are |found only near the southern terminus of the Red Sea, j Craig suggested that the source waters are in this area. |Ross (1972) and Manheim (197*0 have sharply criticized this proposal, however, because it requires subsurface brine transport of 400 to 900 kilometers at an extremely rapid rate across several cross-cutting transform faults :and against reverse density gradients. Manheim (197*0, like White (197*0# favors a water I ;recharge area at the flanks of the Red Sea adjacent to the brine area. He suggests aquifer flow beneath confin ing Miocene evaporites, talcing many thousands of years to ,reach the brine area. Although the present source waters 1 of this recharge area are isotopically dissimilar from the brines, it is difficult to estimate the isotopic composi tion of waters which must have entered the system several thousand years ago. It appears now that the inflowing brine may be considerably hotter than the observed 60° C pool. If this • is the case, then it becomes difficult to explain a process whereby the brine does not undergo isotopic 'exchange. It is apparent that the brine has interacted 1 with evaporites to achieve high salinity and with sili cates (dissolved silica is fifteen times normal sea water). Conceptually, the degree of isotope exchange is a kh | function of brine temperature, water-rock ratio, and i i quantity of brine flow, Clayton (1959) has shown experi mentally that isotopic equilibrium is attained between , water and carbonate minerals within one month at 190° C. . Hydrothermally altered natural calcites are found to i equilibrate with waters at temperatures as low as 70° C# ; Apparently, isotopic exchange is much slower with silicates, Craig (19^9)* however, states that "oxygen [isotope shift , , , takes place in times of weeks for carbonates and feldspars at 100° C," Considering these observations, it seems certain that - the Red Sea brine must equilibrate isotopically with the : rocks it passes through, especially if it attains tempera tures of about 200° C, Nonetheless, the best explanation for the isotopic composition of the brines is recirculated i Red Sea water, largely unaltered by isotopic exchange along its pathway, A possible explanation for this ap- : parent enigma is that the volume of sea water that has passed through the system greatly exceeds the volume of , rock contacted, and the rocks have long ago isotopically ; equilibrated with sea water. i Lacking isotopic data on the rocks of the geothermal ! channelways (or, indeed, any data), we must turn to analogous geothermal systems to support this hypothesis, 1 o ! White (1968) has summarized data on the 8 D- 6 0 relationships in active geothermal systems and oilfield i formation waters (Fig. 9) • For most systems the : previously described "oxygen isotope shift” away from 18 c f meteoric S 0 ratios has occurred. o D variations 1 generally are minor, due to the low hydrogen concentration [ in most rocks. I The Salton Sea geothermal system of the Imperial Valley of southern California represents one of the most ! completely studied systems of this type (White, 1968). | Craig (1966) described the genetic processes involved. 1 ‘ Local meteoric waters percolate into the subsurface along 1 the eastern flanks of the Salton Trough, attain high salinity by interaction with Cenozoic evaporites, are ! |heated, and undergo a large isotopic shift through ex- , change with sediments of the ancestral Colorado delta. Clayton at al. (1968) have examined the oxygen isotope ratios in rocks (carbonates and silicates) from one geothermal well in the Imperial Valley. They have f documented the reverse isotopic shift expected in the ;rocks, thus verifying the exchange process postulated by Craig (1963)* Using this isotopic data, Clayton ejb al. :(1968) are able to calculate an "apparent” water/rock !weight ratio of 0.45* for a batch mixing model of exchange Jbetween the reservoir sediments and the original meteoric :waters• While the Salton Sea is a good example of waters which have been isotopically altered, the geothermal waters 46 j at Wairakei, New Zealand (Fig:. 9) show no shift at all j from the meteoric water line. These dilute thermal waters |attain temperatures of about 270° C as compared with 300° C in the Salton Sea system. As mentioned above, kinetics of ’oxygen isotopic exchange are rapid at these temperatures. j 18 | Clayton and Steiner (1975) have examined 6 0 ratios I in volcanic rocks from Wairakei geothermal wells and have observed extensive equilibration. They concluded that 1 iwaters are flushing through the system rapidly enough to pre-equilibrate the rocks. Calculations, as above, of ; water/rock ratio, give a lower limit of h.3» almost an 1 order of magnitude higher than in the Salton Sea area. Apparently the Red Sea provides an example of an extremely saline brine which is similarly flushing through 1 the geothermal system at a high enough rate to completely ' eliminate "isotopic shift" effects. Important results of these conclusions are: (l) geothermal flow in the Red Sea must occur in well-defined channels, and (2) the brine may have been at moderately high temperature (200° C - 300° C). ; As pointed out by White (1968), high flow rates and water/ i ; rock ratios in the Red Sea system suggest that water | volumes many times in excess of the present brine pools 1 must have passed through the system. This interpretation is consistent with suggested brine overflows at various times in the past. Regardless of the exact brine source and input ____________ k7 itemperature, it is important to know whether the present ;brine pool is the result of significant admixture of i J superjacent Red. Sea deep water* This is especially |important with respect to sulfate in the brine pool which has exactly the same sulfur isotope ratio (+20*3 °/°°) a-s !the overlying Red Sea deep water* ! Craig (1969) has calculated that less than 5 percent ! of the brine is admixed sea water. His methods include an I Q 1 estimate of NaCl saturation at 60 C and mixing-mass I I ‘balance calculations based on F and N0« depletion in the i brine. Obviously the halite solubility estimate is no i i longer meaningful if the incoming brine is considerably 1 1 !hotter than 60° C* Also, it should be pointed out that calculations involving trace element admixture are not 1 1 valid unless it can be established that the aqueous :species considered behave conservatively across the mixing boundary* For example, nitrate reduction may be going on ;at or below the brine-sea water interface due to anoxic conditions. As suggested by Craig (1969), detailed trace ;element profiles of both brine and particulate matter would contribute greatly to our knowledge of the mixing— jprecipitation processes taking place. 1 | Xn summary, it is likely that the brine is composed I I of recirculated Red Sea waters* The present brine i (Atlantis XX Deep) probably is quite close in composition 1 ; to the inflowing brine, although more trace element data i 48 | on the brine is needed to firmly establish, this* Finally, | it is likely that the initial brine is considerably hotter than the 60° C observed in the brine pool and present ' models of brine evolution should be adjusted accordingly* Origin of Cold Brines Oxygen and hydrogen isotope ratios (expressed in the JLS o conventional manner as 6 0 and S D /oo) are plotted in ,Figure 8 for the cold brines and interstitial waters. Ad- 'ditional data on the hot brines and other local waters ,also are included. As can be seen, the interstitial waters from sites 225 and 228 are all isotopically shifted away from the Red Sea water line toward the line represent ing Northeast African meteoric waters. Waters from deeper sections are, in general, shifted closer to the meteoric water line. It is apparent that there is a significant component of isotopically light meteoric waters and, as suggested by Manheim (197^)* they may actually be connate waters trapped at the time of deposition. If so, the isotopic ratios provide significant information about i conditions in the Red Sea basin over the last four or five ;million years. It is unlikely that these pore waters are a result of subsurface meteoric waters flowing in aquifers from the flanks of the Red Sea. Manheim ej f c al. (197^) have shown that chlorinity—depth relationships above the evaporite i l±9 sequences exhibit classic diffusional gradients* Any significant fresh water input would disturb such diffusion profiles which take thousands of years to form. Ap parently, meteoric waters were a significant component of Red Sea waters during and following the dessicated condi tions in uppermost Miocene time. Faber and Schoell (1975) discussed the diverse origins of the cold brine pools. Brines of the Wando and Valdivia deeps are very similar to present Red Sea deep water. Compositional data is not available for the Wando brine, with a chlorinity of 73 °/oo (Backer and Schoell, 1972), but it is almost certainly related to the hot brine system. Like the Atlantis IX 50° C transitional water, this water has been greatly diluted by Red Sea deep water, as its isotopic composition shows. Compositional restraints (Fig. 7) discount any con nection between the Valdivia deep and the hot brines. Also, Valdivia Deep is an atectonic depression with a brine level 400 m more above the Atlantis II Deep. The most likely origin of this brine is solution of evaporites exposed in the Valdivia Deep, although such an evaporite outcrop is unverified. Nereus, Port Sudan, and Erba deep brines (Fig. 8) plot precisely on the predicted isotope line for ocean waters during stages of maximum glaciation. Faber and Schoell (1975) accordingly suggest that these are paleo- _________50 I -------------------------------------------------------------------- i | waters stored in aquifers since the last glaciation, an i explanation consistent with, the theory that transform faults provide conduits for brine discharge and that ris- ; ing sea levels following Pleistocene evaporative cycles is | the primary driving force for the cold brines (Bignell, I 1975? Degens and Ross, 1970)# Isotopic ratios of Suakin and Kebrit brines are not amenable to simple interpretation* However, these ratios I 1 do suggest a mixture of Red Sea water and meteoric water, i similar to the pore water from Deep Sea Drilling Project sites 225 and 228* ! Metalliferous Sediments 1 , Sediments from the hot brine area were first de- I scribed by Miller at al. (1966) and have since been studied : in considerable detail by Bischoff (1969), Hackett and ■ Bischoff (l973)9 and Backer and Richter (1973)# Atlantis II Deep sediments were exhaustively sampled during the : Wando River (1969) and Valdivia (l971» 1972) cruises. | In terms of economic potential, the sediments of 1 i I Atlantis II Deep are singularly important. Bischoff and j Manheim (1969) and Hackett and Bischoff (1973) have | estimated in situ value at about $2.3 billion (1972 ” 1 _ I | prices). Total tonnages of Zn and Cu are 3*2 and 0.8 ■ million tons, respectively. Detailed description of the metalliferous sediments i 1 51 by Bishoff (1969) established the essential nature of* the Atlantis II deposits. The sediments contain up to 95 per cent interstitial brine, are generally very fine-grained (K 62/a), and often are finely laminated. Considering these facts, there is little doubt that these sediments are chemically precipitated from the overlying brines. Ac cumulation rates as established by Ku (1969) and Ku e_t al. (1969) are on the order of 50 cm/l,000 years in the Atlantis II Deep. For comparison, sediments outside the deeps average about 15 cm/l,000 years. Bishoff (1969) defined seven sedimentary facies which are generally correlative throughout the deposit. Facies are defined mainly on the basis of mineralogy and are dif ferentiated as follows: detrital-biogenic, goethite- amorphous, iron-montmorillonite, sulfide, manganosiderite, anhydrite, and manganite. These facies are usually not mutually exclusive and generalizations are necessary in classification. Hackett and Bishoff (1973) have provided a highly schematic diagram (bottom topography is removed) of facies distribution in the Atlantis II Deep (Fig. 10). Minerals identified are listed in Table V and average chemical compositions are in Table VI. Briefly, the youngest and presently depositing facies is the iron-montmorillonite, which has as its major component a ferroan end member of the nontronite group (Bishoff, 1972). The next youngest facies is goethite- _________________________________________________________ 52 Figure 10. Schematic diagram of facies relationships in Atlantis IX Deep sediments (Hackett and Bischoff*, 1973). N S VERTICAL SCALE MONTMORILLONITE □ GOETHITE O il SULFIDE M CARBONATE S 3 BASALT □ (M E TER S) HORIZONTAL SCALE SECTION SHOWING MINERALOGIC FACIES RELATIONSHIPS IN THE ATLANTIS H DEEP. RED SEA Table V. Mineralogy of Atlantis IX deposits Facies Name Mineralogy Montmorillonite ferroan nontronite; amorphous (limonite?) Goethite goethite, limonite, hematite, magnetite Sulfide sphalerite, pyrite, chalcopyrite Manganite manganite, todorokite, woodruffite Anhydrite anhydrite Manganosiderite siderite - rhodochrosite solid solution Detrital biogenic carbonates 55 Table VI# Average composition of Atlantis II sediments (Bischoff, 1969) Detrital Fe Mont- morillonite Goethite- Amorphous Sulfide Manga nite* Si02 27.3 24.4 8.7 24.7 7.5 A12°3 8.4 1.7 1.1 1.5 0.7 Fe^O^ (total) 6.3 37.1 64.2 24.3 30.5 FeO i.1 * 11.7 2.7 13.4 0.4 Mn304 0.6 2.1 l.l l.l 35.5 CaO 23.6 4.8 3.4 2.5 2.9 ZnO 0.08 3.2 0.7 12.2 1.4 CuO .01 0 • 00 0.3 4.5 0.1 0 0 ! \ > 23.1 8.6 3.6 5.7 2.2 S 0.3 3.9 0.6 16.8 0.6 * Based on only 2 \ J X o\ analyses. amorphous, composed mainly of* iron oxides ranging from amorphous Fe(OH)^ . weH “ "crys" ta^1:* - ze<a magnetite. Other major facies are sulfide and detrital. Detrital-biogenic carbonates are the basal facies, below massive (up to 3 m) sulfide layers. The sulfide facies are economically most important and are mainly composed of pyrite, chalcopyrite, and sphalerite. Sulfide facies im mediately overlie the basal carbonates, are thickest in the west and southwest basin, and can sometimes be delineated into three separate units. Xn the Atlantis IX Deep, manganosiderite, manganite, and anhydrite are relatively minor facies. Manganite is more important in Discovery and Chain Deep sediments, where conditions are more oxidizing and dissolved iron has previously been removed from solution, thus favoring manganese oxide precipitation. Manganosiderite is quite common in the Atlantis II Deep but likely anhydrite is of local importance only. Manganosiderite beds are often thin and discontinuous and occur most often within the iron-montmorillonite or iron-oxide facies. Anhydrite is randomly distributed as nodules and layers up to several centimeters thick, suggesting mechanical emplacement or discontinuous processes. 57 Facies Differentiation and Origin From the striking changes in chemical composition and mineralogy of the various facies, it is obvious that conditions in the deep must have been considerably dif ferent at various times in the past* It is even possible, but not necessary, that the chemistry of the inflowing brine was different than the present brine* The intercalated and mixed nature of the various facies suggests strongly that at times in the past there may have been more than one active brine vent. If this is the case, then it might be possible to trace certain facies components to their source areas by stratigraphic correlation* Backer and Richter (1973) have attempted to do so using generalized lithostratigraphic units (Fig. ll) rather than mineralogical facies. The oldest sediments of the Atlantis II, the DOP unit, immediately overlie basal basalts and are generally older than 12,000 years. This unit contains mainly carbonate facies with admixtures of limonite and occa sional lithified pteropods zones ascribed to Pleistocene evaporative cycles. Sulfides are mainly pyrite, but copper content may be locally high* At this time, condi tions must have been much more oxidizing than at present in the brine pool (if indeed a well stratified pool existed)• The limonite content of deep cores in the West 58 Figure 11* Generalized stratigraphic sequence in the Atlantis XX Deep (Backer and Richter, 1973)* UTHOSTRATIGRAPHIC UNITS REMARKS SW-BASIN S AM Sunidic - Amorphous - Silicotic Zone Sulfide Layer SL 0 - >8 m Oxidic- nhydritic 2-5m SO AN Sutfidic-Oxidic- Anhydritic Zone A TLANTIS II D E E P , GENERAL AM Amorphous-Silicatic Zone LP Lepidacracite layer su2 Upper Sulfidic Zone 3 - c o Central Oxidic Zone ZS-lm SUi Lower Sulfidic Zone © DOP ~ 10000 y — UPPER OOP D e trita l- Oxidic - Pyritic Zone ~ 2 ooooy '£jzk. LOWER OOP 25 OCOy +) : - h locally ij a abundant anhydrite W -Passage A $ W - B asin , E - Basin a locally abundant anhydrite CO tacatty detrital and with very reduced thickness € > W-Basin , N-Basin / 0 - 2.5 m / COS / Central Oxidic- / Silicatic Zane carbonates 1 A aV a) a n h y d rite in hill positions E -Basin 153 E BASE HOLOCENE Predominant limanite N-Basin N-Passage and Atlantis Terrace ■ + • basalt flow (+ ) basalt fragments breccia amorpnous facies sulfide facies silicate facies manganite facies limanite and lim anite.hem atite f. detrital - microcaquina facies tfj slumping ■ lithified carbonates © abundant -diatoms □ halite phantoms £ fissures BACKER & RICHTER 1973 60 Basin increases to the north, suggesting a brine vent in the North Basin* The SU^ or Lower Sulfide zone represents the first stable period of* base metal sulfide precipitation* A profile from the Southeast passage outward across the Southeast rim suggests the existence of a stratified brine and a much higher brine surface than at present* SU^ is well developed in the Southeast passage but pinches out at about 1,930 m depth where manganite and limonite facies predominate* Presumably these oxidized phases were precipitated, as today, from an upper transi tional brine layer. Above the 1,905 hi level most traces of hydrothermal activity have disappeared. It is remark able that brine level at this time was nearly 100 m higher than at present. The SU^ discharge zone may have been in the West Basin, where these sulfides appear to be thickest. The Central Oxidic Zone appears to be a chaotic unit with brine source in the West Basin* Sulfides are absent and there is much evidence of slumping. Outside of the West Basin, detrital materials constitute a major portion of this unit. These features, taken together, suggest that the brine pool was flushed out at this time, possibly due to cessation of brine discharge o’ r tectonic disturbance. The Upper Sulfide Unit (SU^) is similar to SU^, 6l indicating that brine stability had been restored. This unit extends out of* the present brine area to the south east to a depth of* about 1970 again indicating a significantly higher brine surface than at present. SU^ can be found in the North Basin also indicating a large stable brine body. The brine source is obscure. The uppermost lithostratigraphic unit is the Amorphous-Silicatic Zone (AM) which is actively depositing today. The AM unit is almost always separated from the SU^ zone by a thin, continuous lepidocrocite layer. Brine source, as delineated by Schoell and Hartmann (1973), is in the Southwest Basin. Recent facies development in the Southwest Basin appears to be unique for the Atlantis XX Deep (Fig. ll). The presence of recent brine activity and several basalt flows has severely disrupted the sequence above the CO zone. Much anhydrite and hematitic material has been emplaced by slumping and turbidites, respectively. Partially open fissures with anhydrite infillings are often observed. These disturbed units have been classified separately as Sulfidic-Oxidic—Anhydrite and Oxidic- Anhydrite. Above these units is the SAM zone which is similar and correlative to the AM zone but with notably enriched base-metal sulfides. The history of deposition is obviously very complex and some of Backer and Richter*s conclusions are specula- 62 tive. Nonetheless, considerable progress has been made and may be summarized as follows: from 25>000 to about 139000 years B.P. the brine vent was in the North Basin, conditions were oxidizing and stratified brines were poorly developed. Base metal sulfide zones were precipi tated from stable, very reducing brines during two dif ferent periods at roughly ^,000 and 8,000 years B.P. Dur ing both of these periods, the brine level was consider ably higher than at the present time. The two sulfide zones are separated by an oxidized unit (CO) apparently caused by disruption of brine stability. All three of these units (SU^, CO, and SU^) probably had their source in the West Basin. During SU^ time the Southwest Basin underwent considerable disrup tion by volcanic processes. The present brine vent is in the Southwest Basin and araorphous-silicatic sediments are presently depositing below the 60° brine. Sulfur Isotopes The determination of sulfur isotope ratios can provide significant information concerning the genesis of sulfur-bearing minerals (Jensen, 19&7 5 Rye and Ohmoto, 197*0* This is particularly true for hydrothermal deposits where source rocks and transport processes are often obscured. Xt is useful to examine general distribu- 63 tions of* sulfur isotopes in nature (Fig. 12) which indicates, particularly, the sulfur isotope ratios of* present-day sea water (+20 °/oo ^ l), hydrothermal sulfides (-15 to +35 °/oo), sedimentary sulfides (-45 to +45 °/oo), and mafic igneous sulfides (—1 to +9 °/oo). Thus, marine sedimentary sulfides often exhibit a large range of isotopic ratios and are considerably depleted ( ^ 45 °/oo) relative to contemporaneous sea water (Kaplan, et al.. 1 9 6 3)# Hydrothermal sulfides in a given deposit r 34 often have a fairly narrow Q S range compared to sedimentary deposits. Hartmann and Nielsen (1966) and Kaplan et al. (1 9 6 9) have examined sulfur isotope composition of various com ponents in the Red Sea system (Fig. 13)« Within the Atlantis XI Deep, sulfates in the geothermal brine (hydro cast sulfate) and interstitial waters are essentially identical to sea water sulfate. Sulfates in the sediment show two ranges around +20 °/oo and 0 to + 10 °/oo. The latter sulfur isotope ratios are apparently due to sulfide oxidation in the cores during prolonged storage. The other group of sedimentary sulfates (<~» + 20 °/oo) may be anhydrite precipitated directly from the brine. It is significant, however, that three anhydrite samples hand- picked (Kaplan ejt al. , 19^9) rather than chemically separated give ratios of +2317» +22.2, and +23.5 °/oo, which are strikingly similar to ratios measured by Shanks 64 Figure 12* Sulfur isotope ratio variations some important natural compounds (Holser and Kaplan, 1966). METEORITE S 2 ' MAFIC IGNEOUS S 2 “ GRANITIC IGNEOUS S O ^ '.S 2 ' VOLCANIC " " . „ S 2 ' PRESENT SEA WATER S O | ' MARINE EVAPORITE S O | ' ATMOSPHERIC S O f' PRESENT FRESH WATER S O | PRESENT MARINE MUD S2 ' SEDIMENTARY ROCK S2 ~ BIOGENIC NATIVE S 3^5627 Figure 13* Sulfur isotope distribution in Atlantis II Deep brines and sedi raents (Kaplan et al., 1969). Total Sulfur ,------------ r - « ------- ,------------ r * ------- 1------- ■ ^ l > ^ f " " ------- 1 Hydrocast Sulfur as Sulfate T ----1 ----1 ----1 ----1 ----1 ----1 I I 1 Interstitial W ater Dissolved Sulfate Sulfur in Sediment as Sulfate ( ------- ! ------- 1 ■ ~ T " ---r— 1 ■ » t — —r .... r * --r * r' r • t Sulfur in Sediment as Sulfide Sulfur in Sediment as Elemental Sulfur r ■ -T — r - i ■■ - t — ■ f - | i r* — Sulfur in Sediment as Separated Sulfide Minerals 8 s3 4 %„ 68 et al. (1974) in adjacent Miocene evaporites. Metal sulfides (and associated elemental sulfur) from the Atlantis XI Deep range from +3*1 to +9*8, averaging +5#7 o / /oo. A number of salient features concerning sulfide deposition in the Atlantis XI Deep should be kept in mind in the following discussion* First, Truper (1969) an<i Watson and Waterbury (1969) have demonstrated that both the Atlantis II Deep brine and sediments are essentially void of bacteria due to prohibitive temperature and salinity. Bacteria may be associated, however, with the suspended matter (Ryan jet al*, 1969) above the transi tional (50° C) brine layer. Second, base metal sulfides with 6 S ratios of about +6 /oo are known to be presently depositing as significant minor components in the uppermost iron-montmorillonite facies (Bischoff, 19695 Hartmann and Nielsen, 1966; Kaplan ejb al., 1969)* Hartmann (l970f 1973) has confirmed the presence of sulfides in precipitating particulate matter. Third, sulfide deposition was the dominant process at various times during the history of the geothermal system. Fourth, Miocene evaporites are believed to outcrop in the Atlantis II Deep, at least in the east bank of the South west Basin (Ross at al. , 1973)* Several hypotheses have been advanced to explain sulfide precipitation in the Red Sea, none of which is 69 entirely satisfactory* Watson and Waterbury (1969) sug gest bacterial reduction of sea water sulfate at the brine-sea water interface* Craig (19&9) kas criticized this theory because the hydrographic profiles analyzed by Kaplan et, al. (1969) failed to show the sea water sulfate 34 necessarily enriched in S above the brine interface. Also, Vinogradov (1962) has shown that bacterial reduction of sulfate within the water column of the Black Sea pro duces negative (-25 to -40 °/oo) sulfur isotope ratios. Kaplan et al. (1969) proposed that the sulfides are formed by inorganic reduction of evaporite-derived sulfate in contact with organic-rich shales at high temperature. Craig (1969) has criticized this suggestion mainly because the high temperatures required are contrary to his 18 interpretation of SO - SO relationships in the brine. Craig provides a third alternative whereby metals are carried in by the brine and H^S is introduced from sedi ments at the depositional site. Further, Craig proposes that /°° & S range is produced abiogenically and at low temperature (60° C), with bacterial growth being prohibited by high salinity. Such low-temperature inorganic sulfate reduction has never been observed either in nature or in the laboratory. 70 ANALYTICAL PROCEDURES General Statement Fifty sediment samples of* sulfide facies material were carefully selected from 13 piston cores from the Atlantis II Deep (Table VII). Eleven of these cores were obtained in 1969 during the Wando River cruise and have been previously described in detail by Hackett (1972). In addition, 35 sediment samples from 21 cores (Table VIII) were provided by Preussag AG, Hannover, West Germany, These samples, most of which were located in cold brine deeps outside the Atlantis II area, were originally collected during the Valdivia (1971* 1973) cruises. Samples were specially selected by Preussag personnel to provide representative examples of possible sulfide mineralization present in the cold brine deeps. The Wando River cores have been stored with their interstitial solution at room temperature since col lection in 1969* Storage has resulted in varying stages of dessication and consequent salt efflorescence. Because of the high salinity and high proportions of interstitial brine (60 to 95 percent by weight), it is necessary to make weight corrections for precipitated 71 Table VII. Summary of piston core data from the Atlantis II Deep Core Number Core Location Water Depth (m) Superpene tration (m) Core Length (m) N. Lat. E. Lon#. 2 21°22.0' 38°03.9' 2131 13.20 17.93 3 21°24.9' 38°04.5' 2027 3.18 8.85 6 21°23.2' 28°03.0' 2078 13.27 8.92 8 21°25.9' 38°04.7' 2069 3.13 8.05 12 21°24.6' 38°03.8' 2078 4.59 5.65 13 21°22.2' 38°02.9' 2144 10.67 11.93 17 21°20.0' 38°04.7' 2164 5.18 7.73 19 21°23.6' 38°03.2' 2070 0 12.14 20 21°23.5' 38°05.3' 2070 3.05 12.21 27 21°22.1' 38°05.0' 2016 4.54 11.87 28 21°21.4' 38°06.2' 2036 5.27 11.75 52KH 21°20.6' 38°05.l' 2181 - - 578PC 21°21.1» 38°o4.9' 2154 — — Table VIII. Summary of* core data for areas outside the Atlantis II Deep Deep Core Number Core Location Water Depth (m) N. Lat. E. Long:. Kebrit 508K 24o43.40* 36°l6.70* 1550 Gypsum 502PT 520P 24°42.13' 24°42.70' 36°24.83 * 36°24.72 * 1196 1191 Vema 525P 23°55.50* 36°27.63 * 1517 Nereus 329P 486K 23°11.63 * 23 11.60* 37°l4.93' 37 15.03* 2447 2430 Thetis 618K 536K 334P 22°38.0 * 22°47.96 * 37^36.10* 37 35.99' 1937 1822 Hadarba—Hatiba 611PT 22°00.34 * 37°52.96» 2263 Valdivia 451K 21°20.30 * 37°57•10 * 1663 Wando 8K 10K 21°21.4 * 21°20.8 * 38°02.3* 38 03.2 * 2001 1991 Albatross 388K 21°12.0 * 38°06.8 * 2133 Shagara 634K 21°07.80* 38°03.23» 2494 Suakin 32 9K 332K 334k 112K 19°36.751 19°36.95* 19 39.12' 19°37.2* 38°43.64* 38^43.95' 38 43.12* 38°47.4* 2838 2848 2777 2787 73 salts• Bischoff (1969) removed brine-derived salts prior to chemical analysis by rinsing sediments with cold de oxygenated, de-ionized water. This technique was rejected in the present study because precipitated salts could be insoluble and original sediment components could be dis solved in some cases. For example, fine-grained sulfides are known to oxidize readily in air and are extremely soluble as sulfates. The technique employed in this study, for Wando River samples, was to oven-dry all sediments at 50° C overnight. One or two grams of dry sediment were then leached with 150-200 mis of cold 0.1 N acetic acid for 2k to k8 hours. The samples were filtered ( 0.45/*) and the acetic acid solution was analyzed for Cl*" concentration. Weight of salts in the original sample was determined using a Gl/total salts ratio of 0.397 based on the brine analyses of Brewer and Spencer (1969)# Valdivia samples had been refrigerated continuously since the time of collection. Samples shipped to the University of Southern California were, of course, at room temperature for several weeks. Sediment samples with interstitial brine were packaged in paraffin-sealed jars or heat-sealed in double plastic bags for shipping. Upon receipt at the University of Southern California, pore waters were removed by squeezing or, in some cases, by 74 washing with de-ionized water. These methods were chosen because the samples showed no signs of dessication or oxidation. Moreover, it was impractical to make salt corrections as previously described in cases where brine composition was incompletely known. Mineral Separations For sulfur isotope analysis, it is extremely impor tant that clean mineral separates, especially sulfates and sulfides, be obtained. Because the fine-grained nature of these sediments precluded physical separations, separations were attempted by chemical means. All samples were first leached with O.IN acetic acid to remove residual pore water sulfates and other soluble sulfates. Anhydrite, gypsum, or bassanite, if detected by X-ray diffraction, were then removed by heating in 6N HC1. If it was desired to analyze sulfur isotopes in CaSO^ minerals, the HC1 reaction was carried out while purging with nitrogen gas to remove H^S generated by re active monosulfides. Sulfate was recovered from solution by BaSO^ precipitation. Residual sulfide minerals, and any elemental sulfur present, were oxidized to sulfate by boiling in aqua- regia-bromine solution. Sulfate was then quantitatively precipitated as BaSO^ by addition of excess 10 percent BaCl#2H20 solution. _________________________________________________________________75 Residual barite was detected by X-ray diffraction in a number of the residues from the aqua-regia leach. Generally, the barite was intimately mixed with various silicates. Prior to sulfur isotope analysis, barite was separated from the silicates by sodium carbonate fusion and reprecipitated as pure BaSO^. A concerted effort was made to separate various sulfide minerals in order to assess possible sulfur isotope fractionations. Such fractionations are important for paleotemperature determinations and the Red Sea provides a "natural laboratory” to check experimentally derived curves. Unfortunately, separation attempts proved impractical because of the extremely fine grain size of these sulfides and extensive alteration in dessicated samples. X-ray Diffraction Mineralogical analysis was carried out on a Phillips X—ray diffractometer unit using CuK^ radiation. A curved crystal monochromator was utilized to discriminate against high levels of secondary iron fluorescence. Samples were mounted for diffraction by filtering ultra-sonically disaggregated sediments onto silver micro- porous (0.45A1) filters. This resulted in fairly well- oriented samples and clay mineral peaks were clearly detectable. 76 X-ray analysis of bulk (dried) samples showed an abundance of salts (halite, gypsum) derived from inter stitial brines and strong reflections from detrital carbonates which in some cases masked the hydrothermal minerals# Therefore, samples were routinely leached with G#1N acetic acid prior to diffraction analysis# This procedure undoubtedly removed some soluble minerals and amorphous material but considerably enhanced the sensi tivity for minor phases# Hydrochloric acid and aqua-regia residues were also routinely analyzed by X-ray dif fraction# In all, this procedure resulted in the accumula tion of considerable detail on insoluble minor phases present# Additionally, it provided assurance that only "clean” samples were used for sulfur isotope analysis. Xn order to evaluate effects of sample storage, X-ray mineralogy was determined on two sulfide facies samples collected and preserved during Chain cruise 6l in 1966# Both samples had been removed from core 127P soon after collection in 1966# Sample 127P--610, collected by J# L. Bischoff, was immediately leached of its salts with de-ionized water, dried, and stored at room temperature in a glass vial# X-ray diffraction analysis (Bischoff, 1969) revealed that originally the principal mineral present was sphalerite, with subordinate manganosiderite• The second sample, 127P—610 to 650, was continuously refrigerated in its original brine-saturated condition in 77 the laboratory of G. Parks at Stanford University, These samples were re-analyzed in 197^ and representa tive diffractograms are presented in Figure 14. Apparently dry, salt-free storage of 127P-610 has resulted in elimination of the sphalerite diffraction peak. The only minerals present are manganosiderite, chalcopyrite, and pyrite. Sample 127F-6lO to 650, on the other hand, shows a strong sphalerite reflection and minor peaks for pyrite, chalcopyrite, and manganosiderite. Thus, it can be inferred that sphalerite and possibly other sulfide minerals may have been extensively altered during storage of the Wando River samples. Indeed, very little sphalerite or chalcopyrite was detected by X-ray diffraction on these samples. Such alteration limits any attempts at mineral separations to samples which have been refrigerated. There is, however, no evidence that sulfur isotope ratios in residual unaltered sulfide minerals would be affected by sample oxidation. Sulfur Isotope Analysis Sulfur-bearing minerals in the sediments and pore water sulfates were all converted to BaSO^ for purifica tion purposes. BaSO^s were reduced to sulfide by re action with graphite at 1,100° C in a nitrogen atmosphere. Sulfide was recovered as Ag^S, Silver sulfide was con verted to SO^ gas by reaction with cuprous oxide at 850° G 78 Figure l4. X—ray diffraction patterns of sulfide facies minerals. 127P-610 stored at room temperature, 127P-(6IO-650) refrigerated. Instrumental settings 2° 20 per minute, 500 counts per second full scale. SL, sphalerite; PY, pyrite; CP, chalcopyrite; Q, quarts and MS, manganosiderite. 127P-610 PY MS GP PY PY 50 10 30 20 Degrees 26 127P- (610-650) SL CP SL+PY PY GO 50 Degrees 26 counts/sec counts/sec 80 for 10 minutes, 0^, 00^, and H^O gas impurities were routinely separated by vacuum distillation. SO^ yields were checked volumetrieally and averaged close to 95 per cent. Sulfur dioxide was analyzed isotopically using the dual collection, 60°, six-inch radius mass spectrometer 3b at the University of California at Los Angeles. 6 S ratios were determined in the conventional manner by comparison to the meteoritic standard (Canon Diablo Triolite). Overall precision of the sulfur isotope 4* O / analysis is better than -0.5 /oo. Xn some cases, however, replicates varied by as much as 3*0 °/oo. These dif ferences are generally due to inhomogeneity of the initial samples. Chemical Analysis Selected sediment samples were analyzed for S, C, Cu, and Zn content. Sulfur was determined gravimetrically at BaSO^, generally on the aqua-regia leach fraction, which approximates total sulfide sulfur. Total carbon and organic carbon concentrations were determined by the LECO gasometric technique. Cu and Zn were determined by atomic absorption spectrophotometry using the model 503® Perkin-Elmer instrument in Environmental Engineering at the University of Southern California. Care was taken to carefully match 81 standard and sample matrices* Where necessary, a deute rium background corrector was used* Major elements (Na, Ca, Mg, K) were determined on four pore waters from cold brine deeps by W. E* Seyfried at Stanford University, standardizing against Copenhagen sea water* Radiocarbon Determinations Eight sediment samples were sent to Geochron l4 Laboratories in Denver for C age determination. Ages are based on the 5»570-year half-life and are referenced to the year 1950. Errors in ages are estimated from count ing statistics as ^ one standard deviation (Table XX), 82 Table XX. Radiocarbon data Core Depth r.14 » C Age Notes 2 1 2 4 3 - 1273 1 7 ,7 5 0 -8 4 0 Age of* STJ^ zone in West Basin 3 2 8 5 - 2 9 0 1 1 ,0 3 0 1 3 2 0 Upper lithif*ied pteropod zone 8 3 5 - 845 2 4 ,3 0 0 ^ 7 5 0 Basal carbonate 6 1 1 5 - 132 1 2 , 9 7 0 ^ 5 0 0 Age of* central oxidic zone in West Basin 6 3 5 - 642 1 8 , 5 0 0 I 6 2 0 Near the top of* detrital-oxidic- pyritic zone 8 1 0 - 820 2 5 , 7 0 0 ^ 8 5 0 Basal carbonate 8 6 7 5 - 685 1 3 ,9 3 0 1 3 8 0 Upper lithif*ied pteropod zone 12 1 1 0 - 130 1 1 ,9 5 5 -2 0 5 Onset of* SU^ in North Passage 3 3 0 - 340 1 3 ,2 2 0 1 3 5 0 Upper lithified pteropod zone 19 8 0 4 - 8 l4 1 3 , 5 6 o l 4 8 0 Upper lithif*ied pteropod zone 27 9 5 5 - 9 6 0 1 2 ,8 5 0 1 3 2 0 Uppermost carbonate facies 1 1 5 5 - 1 1 6 5 2 0 ,3 9 0 1 8 5 0 Basal carbonate 28 9 5 3 - 958 1 3 ,4 7 0 1 3 3 0 Upper lithified pteropod zone 1 0K 1 8 6 - 192 1 4 ,9 4 0 1 3 7 0 Wando Terrace 634k i i o o rw H C V i O 0 - 3 - m H CM 3 , 7 2 5 I 1 8 O 3 , l 4 o l l 8 0 Shagara 534P 1 5 0 - 1 6 0 6 ,7 0 0 1 2 5 5 Thetis Deep 83 RESULTS Introduction Locations of cores utilized in this study are pres ented in Figure 2 for the Atlantis II Deep area and Figure 15 for the Thetis, Nereus, and Suakin deeps. Locations of other cold brine deeps can be found in Figure 1. Radiocarbon ages, including those reported by Hackett and Bischoff (l973)» are summarized in Table IX. Within the Atlantis II Deep five samples were selected to date the uppermost lithified pteropod layer at the top of the DOP zone. Two samples, 27-(1153-1165) and 6-(8l0-820), were selected to date the earliest DOP zone sediments. Other samples from Atlantis II Deep are to establish a detailed chronology of various lithostratigraphic units as noted in Table IX. Cores 3* 8, 12, and 19 are from northern portions of the deep, 27 and 28 are from the East Basin, and cores 2 and 6 are from the West Basin (Fig. 19)* Average sedi— 103 mentation rates range from 145 to 68 cm/ years and are highest in the southern portion of the West Basin. Lowest rates are in the sill area (North Passage) entering the North Basin. 84 Figure 15. Detailed location maps for Suakin, Nereus, and Thetis deeps. Bathy metry from Valdivia cruises (Backer and Schoell, 1972; and Schoell, 1975* personal communication)* Stipled areas represent approximate extent of brine pools. Depth con tours in meters. 85 N ER EU S - ‘ THETIS ^ATLANTIS I I ^VSUAKIN \ D E E P 618K, ■ ■ C v C\ > THETIS D EEP 0 K m 5 L32K SUAKIN D E E P K m N ER EU S D E E P 0 K m 5 86 Sedimentation rates in the DOP zone of 27 and 28 lO3 cm/ years can be calculated for cores 27 and 6, This rate is probably faster than normal detrital sedimentation in the Red Sea (Ku et al,, 1969) due to admixture of rapidly precipitating hydrothermal sediments. Cores from the Wando, Shagara, and Thetis deeps were dated by radiocarbon techniques to establish the age of important hydrothermal processes there (Table XX), The 14,9^0-year age of the Wando Terrace sediments is reason able for non-hydrothermal sedimentation, giving an ac- 103 cumulation rate of about 13 cm/ years. Apparently, hydrothermal sediments are either very recent or are a minor component here. Sediment age in the Shagara Deep shows an apparent reversal with depth. It is possible that slumping has occurred but it seems likely that some small contamination has affected these very young sediments, Ku et. al, (1969) provide a complete discussion of such problems. Stratigraphic Summary Lithostratigraphic relationships within the Atlantis II Deep are summarized in Figure 16 for the cores used in this study. Core locations are vertically justified to 2,040 m-depth, which represents the top of the 60° C brine. Radiocarbon ages and lithostratigraphic units, as • • defined by Backer and Richter, 1973» are included. In 87 Figure 16. Three-dimensional projection of* Atlantis XX Deep stratigraphy. Core locations are vertically justified to the top of the 60 brine (2040 m contour). Litho- stratigraphic units are from Backer and Richter (1973)* Radiocarbon ages (in years) are from Hackett and Bischoff (1973) and this study. 88 2126 DEPTH 1 1 ,0 3 0 24,300 1 1 ,9 5 5 13,220 13,560 2 7 !? if! 12,970 18,500 25,700, I2.850HB 20,390 13,470 17,750 i , 38°02/ 38o 03' 38°04' 38°05' 38°06' 38°07/ 38° 08' F 7-------7-------7 • 1 1 I 13,930 f 20401 2060 WATER 2080 DEPTH 2100 (m) 2120J O-i 2- 4 - 6- 8- LENGTH io- (m) 12- 14- 16- I8J CORE i i i i i i • i i " 11 1 1 ! ' i t CORE LOCATION SUPERPENETRATION CORED INTERVAL AMORPHOUS-SILICATIC ZONE SULFIDIC ZONE CENTRAL OXIDIC ZONE DETRITAL-OXIDIC-PYRITIC ZONE general, the deepest water and thickest sediment sections are in the West Basin (cores 2, 6, and 13)# Lithified pteropod layers are developed in most cores, but are conspicuously absent in core 2# / \ l 4 Ku et al. (1969) have shown that C ages can be determined on these layers and, therefore, aragonitic cementation must have been peneconteraporaneous. Friedman (1972), Milliman et aJL . (1969)* and Gevirtz and Friedman (1966) have indicated that these lithified layers, which occur throughout the Red Sea, are related to late Pleistocene evaporative cycles. The C"^ date on core 2 (Fig. 17) indicates that the top of the DOP zone is about 20,000 years old here. Lithified pteropod zones in this age range are present in other cores from the Red Sea (Ku ejfc al. , 1969)* thus the absence of these zones provides important information concerning the history of hydrothermal activity in the Atlantis XX Deep. This evidence, plus the presence of metalliferous sediment, indicates that initial hydro thermal activity produced a small, stable hot brine pool in the West Basin which inhibited carbonate lithification and was active throughout the last Pleistocene sea level low. Overall sediment thickness decreases from the West Basin into the North Basin and East Basin. Likewise, the age of the top of the DOP zone decreases as follows: core _________________________________________________________________90 Figure 17* Lith.ostratigraph.ic units of cores 2 and 3 showing geochemical varia tions with depth. Radiocarbon ages (in years) from Hackett and Bischoff (1969) and this study. Sulfur isotope ratios (per mil) for bulk sulfide mineral fraction. Organic carbon, carbonate carbon, and sulfide sulfur in weight percent on a dry salt-free basis. 91 CORE 2 DEPTH ' 1 0 0- 200 300 400 500- 600- 700 800- 900- 1000- 1100 - J200- 1300- 1400- 1500- 1600 1700- D O P 1800--------- 17,750 ■ 4 0 -30 -20 -10 0 +10 +20 2 4 6 8 10 12 'ORG C O R E 3 DEPTH 100“ 200 - 300- 400- 500- 600- 700- 800- DOP 24,300 'ORG 92 2, 20,000 years; core 19* 13,500 years; core 12, 12,000 years; cores 3 and 8, 11,000 years; core 27* 13*000 years# Xt seems likely that brine discharge “ First filled the West Basin, then the North Passage and East Basin, and finally the North Basin# Pteropod lithification proceeded at intervals, as in other areas of the Red Sea, until sedi ments were below the brine surface# This conclusion contradicts the suggestion of Backer and Richter (1973) that the initial brine vent was in the North Basin# In addition, the STJ^ zone in the West Basin (cores 2 and 6) is obviously not synchronous with the zone classified SU^ in the North Basin (cores 12, 3 and 8), in the West Basin (cores 27* 28, and 20), or even with core 19 in the northern West Basin# Obviously, the situation is complex and further detailed age dating is necessary to establish correlations with confidence. An important result of this age dating is the establishment of massive sulfide deposition prior to 13,000 years B#P# In core 2, sulfide deposition took place from about 20,000 to 16,000 years ago# In core 6, it occurred from 18,000 to 14,000 years B,P# This sulfide layer was only present in these two cores and must have been deposited from a small stable brine pool localized in the West Basin# Following this initial base metal sulfide deposition, the situation is confused. It appears, however, that cores _________________________________________________________________ 93 1 9 , 2 7 , and 28 show two periods of* sulfide deposition after 13,000 years ago* Xn the North Basin area, there is only one obvious sulfide zone, which appears to be younger than 11,000 years old* Thus, there may have been three periods of base metal sulfide deposition, with the first event localized in small areas of the West Basin* Geochemistry of Atlantis XX Deep Sediments A number of interesting patterns can be recognized by comparing (Figs* 17—2 1) sulfur isotope ratios determined on bulk sulfide separates, weight percent sulfide sulfur, organic carbon, and carbonate carbon calculated on a total (salt free) sediment basis; and the lithostrati- graphic units of Backer and Richter (1973)* In general, sulfur isotope ratios are quite similar for all sulfide facies samples, ranging from +1 to +14*3 °/oo with a +4*7 °/oo average* In nearly every core, how ever, there are occurrences of sulfides (-20 to -45 °/oo) 32 strongly depleted in S * These sulfides generally occur in thin sulfidic layers overlying lithified pteropod zones, indicating that no stratified brine pool existed at these times and normal marine bacteriogenic sulfide reduction proceeded in the sediment pore waters* All of the bacteriogenic sulfides are in the DOP zone except one sample: 6 (125—235)* This sulfide is above the very old oxidic zone noted in cores 6 and 2 in the West Basin, sug— 94 Figure 18* Litliostratlgrapliie units of cores 6 and 8 showing geochemical varia tions with depth (see Fig* 17)* CORE 6 DEPTH CO 1 0 0 - • 1 2 , 9 7 0 2 0 0 - 3 0 0 - 4 0 0 - SU| 5 0 0 - 6 0 0 - • 1 8 , 5 0 0 7 0 0 - DOP 8 0 0 - cm • 2 5 , 7 0 0 ! - 4 0 - 3 0 * 2 0 - 1 0 O + 1 0 * 2 0 2 8 S 34 %C 4 6 S(#) 8 org 10 12 (■), DEPTH 8 0 0 - j cm CORE 8 • COS - SU, DOP 1 3 , 9 3 0 * i i - - - 1 - - - 1 - - - 1 — : t - - - 1 - - - 1 _ i t i - 4 0 - 3 0 - 2 0 - 1 0 0 - M 0 + 2 0 ss3 4 2 4 6 8 % Cqq W t Corg S(#) 10 12 (■), 96 Figure 19. Lith.ostratigraph.ic units of cores 12 and 17 showing geochemical variations with depth (see Fig. 17). 97 CORE 12 SU| ” *11,955 “ DOP •13,220 - - - i i i i i , .i i ._ i___ i___i ----- -4 0 -3 0 -2 0 -10 S S 3 4 ♦10 +20 2 % C CO 4 6 8 (*),C org 10 12 (■), S (•) DEPTH 100- 200- 300- 400- 500- 600- 700- 800- AM SU CORE 17 CO -L _L ± -4 0 -3 0 -2 0 -10 S S 3 4 O *10+20 2 %c 4 6 8 10 1 2 co3^ > ^org (■)* S C ) 98 Figure 20. Lithostratigraphic units of* cores 19 and 20 showing geochemical variations with depth (see Fig. 1 7). 99 CORE 19 DEPTH 100’ 2 0 0 300 900 500 600 700 800 900 1000 ------ ,------ ,------ ,------ ,— T — ------ 1 -------r -1 - « i SU2 CO - " 00S - \ : iQ O c f — 1— 1— 1 _____________) : 13/560 DOP - 1 1— _.j------ 1 ------ 1 ------: « ■ « ------ 1 ------ l— d 8S^ X03 S (O) D E P T H 100 200 300 900 500 600 700 800 900 1000 CORE 20 -90 -30 -20 -10 0 +10 +20 2 9 6 8 10 12 ORG 1 r t — r t ------ 1 1 1 “1 T ” m - SU2 1 ; & : CO ' 1 1----_ i------1 ----- 1 ------1 ------ 1 1 1 1 1 -90 -30 -20 -10 0 +10 +20 2 9 6 8 10 12 Ss3' * «wv<A. c^O). COz ORG S (O) 100 Figure 21- Lithostratigraphic units of cores 27 and 28 showing geochemical variations with depth (see Fig- 17). CORE 27 D E P T H 100“ 200" 300" ''4 0 0 500- 600* 700-■ 800-' 900- 1000- " 12/850 D O P 20,390 1100- ■ 9 0 -30 -20 -10 0 +10 +20 2 8S34 7 o C c 'ORG S ( ° ) C O R E 28 D E P T H 100- 200 - 500- 600- 700- 800“ 900- 1000- 1100- 10 0 +10 +20 2 9 6 8 10 12 g$39 %Cc03 (^), (□ ) . S (O) 102 gesting that it may be correlative with the DOP zone. Four samples (3-(6?-70); 6-(631-636); 12-(369-373)5 and 19 (797-802)) have sulfur isotope ratios between -0.4l and -7.4 °/oo. All of these samples are adjacent to bacteriogenic pyrite samples and represent a transition zone. They may be simply a physical mixture of bacterial and hydrothermal sulfides, formed during the establishment of a stable brine pool. There do not appear to be any significant differences in the sulfur isotope ratios of the various massive sulfide layers. This indicates stable brine conditions and similar processes of sulfide deposition throughout the history of the Atlantis IX Deep. Within a given sulfide zone, there are subtle fluctuations in sulfur isotope ratio, but no persistent patterns can be recognized. Such fluctuations may be due to admixture of small amounts of bacteriogenic sulfides which settle in from the brine-sea water interface. An interesting result of this study is the discovery of abundant occurrences of barite as a minor component associated with sulfide facies materials. Figure 22 is a summary diagram of sulfur isotope ratios for coexisting barites, anhydrites, and metal sulfides. Barite is en- 34 o / riched in S by an average of 5*5 /°° over coexistxng sulfides, however, the fractionation is quite variable. Xf the barite was formed by an isotopic equilibrium pro— ______________ 103 cess, temperatures of* 700-800° C (Fig* 16) would have been required. Obviously, such temperatures have not been achieved and the most likely origin of barite is in situ oxidation of* sulfides. Anhydrite separates (Fig. 22) range from +20.5 to +22.0 °/oo. Sulfates with sulfur isotope ratios close to +20 °/oo are almost certainly formed by precipitation of anhydrite from recent sea water sulfate. Sample 52KH is from a partially open fissure in hematitic turbidite deposits in the Southwest Basin. Mineralization was mainly anhydrite, with subordinate amounts of metal sul fides intimately associated. This is probably a case where sulfates and sulfides precipitated together, and the fractionation factor of 18.5 °/oo requires temperatures of 350°-400° C for isotopic equilibrium processes (Sakai, 1968) . Anhydrite samples from cores 12 and 578PC, however, have sulfur isotope ratios between +21 and +23 °/oo, similar to isotopic ratios reported by Shanks ejb al. (197*0 in Miocene evaporites flanking the hot brine area. In core 12, anhydrite occurs as a massive layer in the lowermost meter of the core. This layer does not correlate with other nearby cores (Backer and Richter, 1973)* As men tioned earlier, the Southwest Basin in the area of core 578PC is probably disrupted by slumping. Thus, it seems quite probable that both of these occurrences are Miocene 104 Figure 22. Sulfur isotope fractionation between co-existing sulfur-bearing minerals in tbe Atlantis IX Deep. Sample numbers for Wando River and Valdivia cores. 105 I I I I AA—o ---o 2 - ( 1 4 6 7 - 1 4 6 9 ) A O 3 - ( 1 4 5 - 1 5 5 ) A ----- O 3 - ( 2 0 5 - 2 1 5 ) 12 - ( 4 8 5 ) <A-------------------nn 1 7 - ( 3 8 5 - 3 9 0 ) A -------CD A --------------O 1 9 - ( 4 3 5 - 4 5 5 ) A = O 19 - (620640) &> 1 9 - (650660) A ----------- O 19 - (690700) A-------------------□ 52KH - ( 3 1 4 - 3 1 7 ) 5 7 8 P C - ( 1 1 1 - 1 2 1 ) CXI ____________i____ i l ___________i 0 5 1 0 1 5 20 25 8 S 34 o/oo A SULFIDE O BARITE □ ANHYDRITE 106 evaporite material brought in by slumping and mass move ment. Wando, Albatross, and Shagara Deeps Sulfur isotope and chemical data on samples from out side the Atlantis XI Deep are compiled in Table X. Wando, Albatross, and Shagara deeps are discussed together in this section because of proximity and possible connection to the Atlantis XX Deep by hot brine overflow. Core loca tions are given in Figure 2. In the Wando Basin, samples from the upper part of core 10K and from core 8K are typical hydrothermal sul fides. These sediments are similar to the Atlantis II sulfides in sulfur isotope ratios, sulfide mineralogy, and macroscopic properties. The present-day existence of brine in part of the Wando Basin and inferred high-brine levels in the Atlantis II Deep in the past indicate an almost certain connection by brine overflow. Albatross Deep presently contains brine which is 18 chemically and isotopically ( § D- So ) similar to Atlantis II brine but the samples of the bottom (core 338K) are typical detrital sediments. The location (Fig. 2) of this core, however, seems to be on the flank of Albatross Deep and may be poorly placed for sampling hydrothermal sediments. Core 63^K from the Shagara Deep is located about __________________________________ 107 370 m deeper and directly along a topograph.!c low from Albatross Deep# No brine was recovered in Shagara Deep but pore water chlorinity was found to be 113 °/oo (Backer and Schoell, 1972)# Xn the present study, pore waters were washed from samples 634K-(l30-l4o) and 634K-(240-250) and analyzed for Cl, Na+, K+, Mg^+, and 2 + o • Ca « Correction to 113 /00 Cl*” gave the same result for both samples. This brine is presently dissimilar from the Atlantis XX brine or any mixture of Atlantis II brine and sea water (Fig. 5)• Sediments from Albatross Deep are typical biogenic carbonates# However, they contain small concentrations of sulfides which have sulfur isotope ratios analogous to Atlantis II base metal sulfides# Sulfur concentration is very low and it is apparent that bacterial sulfate reduction has been inhibited by the presence of brine over the last 4,000 years or so. The most plausible explanation is that these sulfides are the result of a brine overflow into Albatross Deep and then to Shagara Deep. It is possible that a cold brine existed in Shagara Deep prior to this event, and brine mixing has destroyed the chemical characteristics of the original Atlantis II brine. In general, these isotopic studies of sediments from deeps near the hot brine area indicate that a substantial brine overflow occurred about 4,000-6,000 years ago. 108 Thetis, Suakin, and Gypsum Deeps Samples from Thetis and Suakin deeps were studied in some detail because of obvious hydrothermal influence. Backer and Schoell (1972) describe sedimentary magnetite from the Thetis Deep, similar to that found in the South west Basin of the Atlantis XX Deep. However, no brine has been detected in this basin. Brines and hydrothermal sediments from the Suakin deep have been studied by Baumann ejt al. (1973)* While compositionally distinct from the Atlantis II brine, the Suakin brine does exhibit a small temperature (up to 24.6° C) and chlorinity stratification. Hydrothermal sedi ments occur in the Suakin Northeast Basin, where iron and manganese concentrations range to 29 percent and 11 per cent, respectively. Cores 6l8K and 53^P from Thetis Deep, North Basin (Fig. 15), are by far the highest grade hydrothermal sediments found outside of the hot brine area. Mineral- ogically (Table X), they consist of magnetite or hematite with subordinate amounts of base metal sulfides. With no existing brine, little can be directly concluded about precipitation processes. We can infer, however, that conditions must have been quite similar to present Atlantis XI Deep where magnetite is forming in the Southwest Basin sediments. 109 Table X, Geochemical data from the cold brine deeps INTERSTITIAL WATER _ _ _ _ _ _ _ SEDIMENT MINERALOGY Deep i Core Depth (cm) c 1 4 Age S04=(ppm) s34(so4) 7«S Sulfide 7«C C03 7oC Organic s 34 Major Minor Kebrit 508K 10- 20 6.4? 1,05 -44,5 Ct Q,Ar Gypsum 502PT 60- 86 18,9* - - +21,19 Gy,Bs +20,87 520P 100- 20 18,7* - - +20,07 Gy,Bs +20,11 Vema 525P 31- 36 0,34 - - -33,46 Ct Py Nereus 529P 54- 59 - . . . Ct 486K 11- 18 1034 +15,56 1,58 6,03 1,51 -19,15 Ct W y 60- 65 1157 +13,89 1,66 5.47 1,88 -25.73 Go K,I,M,Py? Thetis 618K 122-132 - - 3,85 - U 7 +12,10 Mg Py»sp +11,21 536K .18- 22 6903 +13,11 1,02 4,87 2.43 -30,16 Ct,Q ¥ , P y 48- 52 6823 -21,45 2.67 3,94 3,94 -34,11 Ct,Q Py,K 148-152 3460 +19,31 0,05 6,73 0,60 - Ct,Q Ar,K 534P 85- 95 3425 +19,21 .03 3,72 0,52 - Hm - 150-160 6700y 2194 - 0,16 3,80 0,47 - Hm K,Mg 190-200 6812 +12,59 2,06 - 0,90 + 9,88 Hm,Mg K,I Hadarba- 611PT 419-433 - ■ 0,13 5,30 0,96 - Ct Q,Ar,D Hatiba Valdivia 451K 14- 22 - - 2,18 5,06 1,94 -17,68 Ct Ar,Q,D,Py Wando 10K 38- 53 4734 - 2,43 2,25 1,64 1,75 - 9,92 Py,Sp Ct,K,Q,D - 0,04 - 0,15 53- 67 6173 + 7,18 0,56 - 0,89 +17,19 M K +14,98 186-192 14,940y 2739 -24,35 7,56 2,16 3,69 -31,25 Q»cr Py,K Table X, Geochemical data from the cold brine deeps Deep Core Depth (cm) c14 Age INTERSTITIAL WATER SEDIMENT MINERALOGY S04=(ppm) S34(S04) IS Sulfide 7oC C ° 3 Organic s34 Major Minor Wando 8K 55- 86 4832 0,53 . 0,72 + 3,25 . (cont'd) Albatross 388K 30- 42 . ■ 0,57 4,76 3,57 . Ct,Q Py 60- 70 606 +34,45 5,29 5,19 3,16 -26,39 Ct Q , K , P y 280-290 915 + 5,62 2,83 4,59 1,54 -25,64 Ct Q,K,I,M,Py Shagara 634K 130-140 3725y 1189 +14,41 0,19 5,96 1,05 - 1,08 ct Ar,D,K,Q,Py 180-184 - - 0,06 7,07 0,85 + 3,35 ct Ar,Py 240-250 3140y 937 +11,88 0,08 7,02 1,34 + 5,90 ct Py,K,D,Q Suakin 329K 610-615 2113 +16,15 0,11 4,70 0,45 . 1 ct Q 332K 280-290 3197 +21,04 0,11 4,73 0,36 - Ct,K,Q I,M 805-815 344 -13,69 2,89 3.63 1,01 -32,00 Ct,Q K,I,M,Py,Ar 334K 51-56 - - 0,16 2,97 2,47 -23,98 Ct,Q Ar 350-360 3208 +21,39 0,11 2,41 0,75 +18,98 Ct,Go Q,I,M,K +15,99 112K 100-120 4274 +11,56 0,75 3,17 1,07 -27,44 Ct,Q M,I,K,Py 111 Sulfur isotope analysis of sediments from Thetis North Basin indicates the presence of hydrothermal sul fides similar to those in the Atlantis II Deep# Obviously, there is no possibility of physical connection with the Atlantis II system# This important discovery, however, indicates that similar geothermal systems can be generated at widely separated intervals along the axial rift zone. There is some evidence of a mid—water echo reflector (M# Schoell, personal communication) in the Thetis Deep South Basin, suggesting the existence of a brine pool. No water samples were obtained, however, and all geochemical evidence indicates normal detrital sediments in core 536K from this basin (Table X)# Samples from the Suakin Deep sediments were specially selected from Cu and Zn enriched zones as analyzed by Baumann jet al# (1973)• Evidence of hydrothermal sulfide deposition is not clear in this basin. Only one sample (33^K-350-36o) contains sulfides which are not typically bacteriogenic# This sample contains abundant goethite and the S S^^ ratio of +16,5 °/oo in the sulfides, plus the low organic carbon content (Table X) suggests hydro- thermal processes# The other example of possible hydrothermal processes is the Gypsum Deep (Fig. l). Again, no brine was detected in this deep but the two sediment samples analyzed are almost pure CaSO^. Sulfur isotope ratios indicate deposi- 112 tion from recent sea water. Other Deeps and Interstitial Waters Other deeps sampled include Kebrit, Vema, Nereus, Hadarba-Hatiba, and Valdivia (Fig. l). Of these, only the goethite-containing sediments of Nereus Deep provide any indication of geothermal activity. All sulfur isotope ratios of sulfide minerals are indicative of normal low— temperature marine processes. Backer and Schoell (197^) reported detectable quantities of H^S in the Kebrit Deep brine, however, gas samples were not available. Surficial sediments from this deep show no hydrothermal influences. Pore waters were squeezed from most of the samples received and sulfur isotopes were determined on dissolved sulfate. These data (Table X) for the most part only indicated the reactive nature of sulfides in the sediment. In most cases, o S ratios are similar to sea water or show evidence of oxidation of sedimentary sulfides. Kaplan et al. (1969) made similar observations. In one instance, (388K-6O-70) pore water sulfate from the Albatross Deep 34 is enriched in S , indicating in situ sulfate reduction. Carbon - Sulfur Relationships Organic carbon and carbonate carbon concentrations are plotted against sulfur isotope ratios in Figures 23 and 24. In general, these diagrams emphasize the separa- 113 Figure 23. Variation of sulfur isotope ratio with organic carbon content of sediments. Sulfur isotopes for bulk sulfide mineral fraction. Organic carbon in weight percent on a dry salt-free basis. % ORGANIC CARBON V. 8 . 0 7 . 0 6.0 5 . 0 4 . 0 3 . 0 2.0 1 . 0 • ATLANTIS n □ SHAGARA O TH ETIS OSUAKIN A WANDO x OTHERS A H Y D R O T H E R M A L ^ DETRITAL . : * □ . * • S » t « 111 II111 I I 1 1 I I 1 1 11 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 1 11 111 1 1 1 1 1 1 111111 11 i 1 1 1 11 - 4 5 - 4 0 - 3 5 - 3 0 - 2 5 - 2 0 - 1 5 - 1 0 - 5 S S 3 4 (°/oo) + 5 + 1 0 + 1 5 115 Figure 24. Variation of* sulfur isotope ratio with carbonate carbon content of sediments (see Fig. 23)* 7.0 r O S'60 < ° 5 .0 UJ 5 24.0 S « 3.0 o >82.0 1 . 0 • ATLANTIS n OTHETIS A WANDO NORMAL DETRITAL □ SHAGARA O SUAKIN x OTHERS o □ HYDROTHERMAL IRON OXIDE CONTAINING A. i » » -1 I i i m I i i i i I i i i i I i i i i I i i » i I i i i i I i i i i I i i I r l t i . J A X i A l i f U i -45 -40 -35 -30 -25 -20 -15 -10 -5 0 +5 -H O +15 8 S 3 4 (%o) 117 tion of sulfides into two distinct groups based on sulfur isotope ratios* Organic carbon contents are overlapping in the two groups (Fig. 23), with detrital sediments generally enriched. Figure 2k shows that detrital carbonates are greatly diluted in areas of hydrothermal sulfide deposition. In some cases, notably in the Shagara sediments, hydrothermal sulfides may be admixed with detrital carbonates. Sample 19 (797-802) contains manganosiderite. Samples with negative & S ratios can also be divided into two groups. Those containing iron oxides are notably depleted in carbonate content due to admixture of hydrothermal sedi ments or dissolution of carbonates. Xt is quite surprising that the organic carbon content of the hydrothermal sediments is so high, consider ing the high sedimentation rates and low content of detrital carbonates. The most likely explanation is that a reactive portion of detrital organic matter in normal marine sediments is decomposed by bacterial activity. In the sediments containing supersaline brines, bacterial activity is inhibited and a large fraction of the detrital organic matter is preserved. This argument is supported by the lack of bacterial sulfate reduction in the hydrothermal sediments. An interesting further study would be a comparison of the organic composition from 118 hydrothermal and normal marine areas of the Red Sea to determine which organic fraction is readily degraded by bacteria. 119 GEOCHEMICAL MODEL General Statement Theoretical evaluation of chemical equilibria in high ionic strength aqueous solutions like the Atlantis II brines is a difficult problem. Two approaches have generally been used to determine solubility relationships with solid phases. One method involves the use of inter action coefficients (Harned coefficients) to determine mean ionic activity coefficients for various electrolytes in solution. This is an empirical procedure which requires much experimental data but it is the only useful method for very concentrated mixtures of strong aqueous electrolytes. Lerman (1967) and Wood (1975) have shown that this so-called "Harned*s Rule” can be useful in evaluating equilibria with very soluble salts in both the natural environment and laboratory experiments. At present, how ever, this empirical approach cannot be extended easily to include high-temperature brines or solubility relation ships of minor components. The second approach, and the one utilized here, is to employ thermodynamic equilibrium constants for aqueous complexes and compute individual ionic activity coef- 120 ficients for each dissolved specie* This is essentially an extension of the model of Garrels and Thompson (1962) which successfully predicts speciation in sea water at 25° C and atmospheric pressure* Helgeson (1969) has compiled an internally consistent set of data for aqueous solutions up to 300° C* Xn general, the data of Helgeson (1969) must be considered as only approximate* Equilibrium constants are often only estimates and much more experimental data is needed* The problem of calculating individual activity coefficients is particularly difficult in solutions of ionic strength greater than 0*5M* However, the present data are the best available for hydrothermal solutions and are very useful in making generalized predictions* Conditions Compositions of the Atlantis XI brines have been discussed earlier (Tables I and II)* The 60° C brine is approximated by a 4*5M NaCl solution with significant amounts of Ca. The 50° C brine has an ionic strength of about 2.5M, due to mixing with Red Sea deep water* Relationships discussed by Craig (1969) seem to indicate that the 60° C brine has not been significantly diluted or mixed with Red Sea deep water* Also, in spite of the notable influx of new brine since 1964, detectable changes in chlorinity have not occurred (Brewer et; al*, 121 1971; Schoell et al. , 1975). Faber and Schoell (1975) have also shown that S o18- £ d relationships in the brine are unchanged, Thus, available evidence suggests that the 60° C brine is compositionally very similar to the inflow ing fluid* Observations of temperature increases in the brine pool, discussed above, indicate that new brine influx must have a temperature of about 200° C* An interesting independent method of estimating initial brine temperature is the Na—K-Ca geothermometer developed by Fournier and Truesdell (1973) For hot springs and other natural waters* Using molar Na, K, and Ca data for the 60° C brine and the empirical equation: l/T^ogfMj^/MjjO+l/S logC^M^/^) (l) one obtains an equilibration temperature of about 150° C. Considerations of silica solubility in the brine provides a similar estimate (Kennedy, 1950)* Xn general, these methods tend to provide a lower limit of brine temperature since the brine may have been considerably hotter but o continued to equilibrate only down to about 150 C* Thus, available evidence indicates that subsurface brine temperature is considerably in excess of the 60° C observed in the brine pool and there is probably little mixing between the brine and overlying sea water* In order to evaluate the important process of ore transport into the Atlantis XX Deep, it seems reasonable to assume that 122 the inflowing brine is considerably hotter but composition- ally quite similar to brine pool. Minor elements are, of course, being precipitated and are more difficult to evaluate. For the purposes of this model the analyzed concentrations of heavy metals (which represent lower limits) are utilized. Oxidation potential of the brine is likely controlled by interaction with solid phases during subsurface move ments, Hackett and Bischoff (1973) provided a detailed discussion of the transformation of hematite to magnetite in the Southwest Basin of the Atlantis IX Deep, They found that for present iron concentrations, pH, and temperature, the Atlantis II brine is close to magnetite- hematite equilibria. Therefore, upper limits of oxygen fugacity (fn ) can be estimated using the equation: 2 Fe203=Fe304+l/2 O., (2) Temperature-dependent equilibrium constants for this re action (Helgeson, 1969) can be used to derive fn condi- 2 tions in the inflowing brine at higher temperatures (Kajiwara, 1973b). The assumption of magnetite-hematite equilibria in the incoming brine is probably reasonable considering that the inferred heat source is recent basic intrusives in the rift zone. Such rocks contain ubiqui tous minor magnetite and any dissolved oxygen in the brine would be removed by production of minute quantities of hematite. 123 An alternative method of* estimating f^ is based on the observed association of* sphalerite and barite in the sulfide facies# The following reaction can be written for equilibrium between these two minerals: BaSOZ|+Zn2+=ZnS+Ba2++ 2 0 ( 3) The equilibrium constant for this reaction, as a function of temperature, can be derived from data in Helgeson (1969). Z inc-barium ionic activity ratios can be calculates using concentrations in the present brine pool and the individual ionic activity coefficients# Results of both oxygen fugacity estimates (Fig# 25) are surprisingly similar and fn conditions are relatively more oxidizing 2 at higher temperature, for equilibrium with either of these mineral pairs# The final parameter needed for this model is an estimate of the pH of the brine# Miller et al# (1966) and Shishkina and Bogoyavlenskiy (1970) have measured pH in the hot brines# Measurements were done potentiometrically on brine samples cooled to 25° C# Values selected for this study are 5*5 for the 6O0 C brine and 5*9 for the 50° C# Adjustment of these pH values to in situ tempera tures will be discussed below# In summary, the assumptions and conditions chosen for this geochemical model of brine evolution are as follows: (l) complexing calculations and data in Helgeson (1969) are useful in evaluating geochemical relationships in hot, 124 Figure 25. Oxygen fugacity estimates for Atlantis XX brine. 12 5 LO G \ -30 -40 -50 -60 - 70- M A G N E T I T E - H E J W I 7 E B A R I T E - S P H A L E R I T E j. 0 50 100 150 200 TEMPERATURE (°C ) 126 saline brines. (2) The inflowing Red Sea brine is considerably hotter but compositionally similar to the present brine pool. The temperature range evaluated here is 25° to 250° C. (3) Oxygen fugacity data can be derived from the equilibria of either magnetite-heraatite or barite-sphalerite in contact with the brine. (U) Measurements of pH at 25° C in the brine are reasonably accurate and in situ values at higher temperatures can be calculated. (5) The unassessed pressure effect on species distribution is relatively small. (6) Chemical equilibrium is attained in the brine. Data Equilibrium constants for aqueous species (Table XX) and selected minerals are from Helgeson (1969) except constants for NaCO^ , MgCO^°, MgHCO^+, and CaHC0^+ which are from Lafon (1969). Extrapolations based on constant enthalpy of reactions were necessary to obtain equilibrium constants for a number of species above 200° C, and at all temperatures higher than 100° C for Lafon*s data. Xn the case of NaSO^“, which is of pivotal importance in sulfate- sulfide equilibria because it is a major sulfate species in the brine, the dissociation constant used is from the experimental work (Fig. 26) of Fisher and Fox (1975) and Styrikovich e_t al. (1968). A noticeable absence from available data is the equilibrium constant for NaHCO^° _______________________________________________________________ 127 Figure 26. Dissociation constants for NaSO^-. 128 L O G KncSoI -4 Fisher and Fox (1975) 50 100 150 200 250 300 350 TEMPERATURE (* t) 129 complex which, was calculated to be a major bicarbonate specie in the Garrels and Thompson (1962) model and might be quite important in the pH calculations (see below)* Individual ionic activity coefficients were calculated from the extended Debye—Huckel limiting law* lQe ^ = jfl b- * 1 + B°x <*> In this equation z is the absolute value of the ionic charge, I is the ionic strength of the solution, a is an ionic size parameter, ^^ 3-s "the activity coefficient, and A and B are Debye-Huckel constants* The term B° I is to correct for high ionic strength effects on^^ and is valid only in the presence of "swamping” concentrations of NaCl (Helgeson, 1964, 1969)* A, B, and B° are compiled by Helgeson for temperature and ionic strength rages useful in this model* Values for a are available only for 2 5° C and for certain species* Temperature effects are considered negligible here and equalities of a are assumed for similar complexes• Activity coefficients for uncharged species are all assumed equal to Y CO^* Values are derived from curves in Figure 27, where I is the true ionic strength* Experi mental data are from Helgeson (1969) and Robinson and Stokes (1959)# 130 Figure 27, CO^ activity coefficients utilized in geocbemical model. Data from Helgeson (1969) and Robinson and Stokes (1959). 131 l £ A C T I V I T Y C O E F F I C I E N T C 0 2 0.8 0.7 0.3 0.2 0.1 O . O l 5 3 2 TRUE IONIC STRENGTH ( I ) 132 Calculations Distribution of dissolved species is calculated using mass balance and mass action equations as first described by Garrels and Thompson (1962). Computations were carried out by computer using an iterative process. Briefly, calculations are as follows. Initially, ionic strength is computed assuming measured concentrations are equal to free ion concentrations. Once ionic strength is computed, 'JSV s can be computed from the Debye-Huckel equa tion, With activity coefficients known, molalities of complex ions are calculated from equilibrium relationships assuming initially that IQN. Molalities of free ions are then calculated by substituting molalities of complexes into mass balance relationships, as follows: ^PREE ION=MTOTAL~ 51 MCOMPLEX IONS (5) The whole cycle is then iterated until there is no further change in speciation. In all, 52 species were evaluated (Table XI), Since the equilibrium constants of complexes change with temperature and many complexes contain H , in situ pH may be quite different from pH in a cooled solution. It was thus necessary to evaluate pH at temperature by a hydrogen mass balance calculation similar to that of Bischoff and Dickson (1975)# Total hydrogen concentration was first evaluated at 25° C according to the following ________________________________________________________________133 Table XI, Dissociation constants of aqueous complexes LOG K Equations 25° 50° 60° 100° 150° 200° 250s + 2- = hT + SO, - 1,99 ■ 2,27 - 2,40 - 2,99 - 3.74 -4,49 -5,41 + '2- = K + SO. - 0,84 ■ 1,00 ■ 1,06 ■ I,'30 v l ^ O -1,94 -2,28 2+ 2- = Ca + SO, ■ 2,31 ■ 2,40 - 2,50 - 2,70 - 3,10 -3,60 ■4,40 2+ 2- = M + SO, 2 4 - 2.25 - 2,60 ■ 2,70 ■ 3,20 I' ■ 3,90 ■4,80 -5,90 °+ 2- = Na + SO, - 1,06 - 1,35 - 1,46 ■ 1,93 - 2,57 -3,30 -4,10 + 2- = 2H + SO, 125,55 113,20 109,31 94,31 79,31 67,13 57,1 + I- = H + SO. 4 132,53 120,0 116,02 100,93 86,03 74,09 64,1 = SO 144.63 133,32 128,72 112,70 96,68 83,74 11,1 + 2- = H + CO, ■10,32 -10,17 -10,15 ■10,16 -10,29 ■10,68 ■11,43 + 2- = H + HCO - 6,35 ■ 6,31 ■ 6,32 - 6,45 - 6,73 -7,08 -7.63 2+ 2- = Ca + CO - 3,20 ■ 3,40 - 3,50 - 3,90 - 4,50 -5.20 -5,90 + 2- = Na + CO ■ 1,27 - 1.81 - 2,03 - 2,89 - 3.95 -5,00 -6,10 2+ 2- = Mg + CO„ ■ 3,40 - 3,42 - 3,44 - 3,55 ■ 3,70 ■3,85 ■4,05 + 2+ 2- = H + Mg + CO ■11,22 -11,70 ■11,93 ■12,93 ' -14,35 -15,7 -17,1 + 2+ 2 = N + Ca + C03 ■11,58 -11,82 -11,95 -12,59 -13,15 -13.7 -14,3 = H+ + Cl" 6,10 5,00 4,56 2,90 1,23 0,06 -0.67 = Na+ + Cl" ■ - - - 0,97 0,42 -0,15 = K+ + Cl" - - - - 0,90 ' 0.30 2+ = Mg + OH ■ 2,60 ■ 2.7 - 2.8 ■ 3.1 1 - 3.6 -4.1 -4.6 = Fe2+ + l/402 + H+ ■ 7.75 - 6,35 - 5.87 - 4,16 - 2.48 -1.15 -0.1 KSO," 4 CaS0,° 4 « g s o 4 ° NaSO." 4 H2S + zo HS" + ZO S= + ZO 2 HCO 3" h2co3° CaC03° NaC03" MgC03? MgHC03+ CaHCO + HC1° NaCl° KC1° MgOH+ Fe3++1/2H20 134 Table XI, Dissociation constants of aqueous complexes (continued) Equations_ _ _ _ _ _ _ _ _ _ _ _ _ _ _ 25° FeOH+ = Fe2+ + 0H“ -5,7 2+ 3+ FeCl = Fe + Cl - 1,48 FeCl2+ = Fe3+ + ZCl“ -2,13 FeCl3° = Fe3+ + 3Cl“ -1,13 FeCl," =Fe3+ + 4Cl" 0,79 Cu2+ + ^ H20= Cu+ + l/402 + H+ -18,1 CuCl2' = Cu+ + 2Cl’ -4,94 CuCl2" = Cu+ + 3Cl" -5,14 CuCl = Cu2+ + Cl" -0,01 CuCl2° = Cu2+ + 2C1" 0,69 CuCl3“ , =Cu2+ + 3Cl' 2,29 CuCl,2' = Cu2+ + 4Cl“ 4,59 ZnCl = Zn2+ + Cl" - 0.43 ZnCl2° = Zn2+ + 2Cl" -0,61 ZnCl3“ = Zn2+ + 3Cl“ -0,53 ZnCl2" = Zn2+ + 4Cl" -0,20 PbCl = Pb2+ + Cl“ -1,60 PbCl2° = Pb2+ + 2Cl“ -1,78 PbCl3“ = Pb2+ + 3Cl“ -1,68 PbCl42" = Pb2+ + 4Cl“ -1,38 _ _ _ _ _ _ _ _ _ _ _ LOG K 50° 60° 100° - 5.6 - 5.5 -5.4 - 1.96 - 2,15 - 2.94 - 2.62 - 2.82 - 3.63 - 1.76 - 2.00, - 3,00 - 0.05 - 0.40 - 1,63 -16,1 -15,35 -12,83 - 4,94 - 4,97 ■ - 5,06 - 5.18 - 5,22 - 5,39 - 0.53 - 0,73 - 1,54 0,06 - 0,18 - 1,15 1,48 ,0.26 - 0,04 3,54 2,00 1,63 - 0,90 - .99 - 1,82 - 1.12 - 1,32 - 2,13 - 1,14 - 1,35 - 2,23 - 0,89 - 1,14 - 2,14 - 1,63 - 1,65 - 1.73 - 1,85 - 1,87 - 2.04 - 1.81 - 1,86 - 2.13 - 1.59 - 1,69 - 2,05 150° 200° 250° 5.5 -5.7 -5.9 3.98 -5.1 -6.2 4,72 -5.9 -7.1 4,30 -5.7 -7.1 3,23 -4.9 -6,6 10,34 -8.35 -6,64 5,35 -5.8 -6,5 5.77 -6.4 -7.2 2,57 -3.7 -4,8 2,36 -3.7 -4.9 1.52 -3.1 -4.6 0,18 -2.0 -3.9 2,78 -3.9 -4.8 1.19 -4,4 -5.5 1.34 -4.8 -6.0 3,35 -5.0 -6.4 1,88 -2.1 -2.4 2.29 -2.6 -3.1 2,50 -3.0 -3.6 2,57 -3.2 -4.0 equation: X o T A l / 5 * W 5 ° C ) + M H S 0 l t - + M H C ( > 3 + 2 M H 2 C 0 3 ( 6 ) +2MH S+MHS +MHCl+MCaHCO + + M,, + 2 3 + MMgQH+ The molality of free hydrogen ion at temperature was then derived using the ^qTAL va- * - ue ^rom 25° C and rearrang ing the hydrogen mass balance equation, MH+(T°r)=MHTOTAL “ ^ ^ i (7) Activity coefficients for H+ are from equation (4). Hydrogen ion concentration contributed from water dis sociation, due to the changes in the equilibrium constant of water with temperature, is insignificant• Speciation Distribution of hydrogen-bearing species for the 60° C Atlantis XX Deep brine composition as a function of temperature is presented in Figure 28, By far the most important species is H^CO^ at these low pHs. In situ hydrogen ion activity continually rises with temperature. This is due mainly to the increased stability of the MgHCO^ complex. Essentially, this transfers carbonate from the H^CO^0 to the HCO^*” "reservoir” and hydrogen ion is released in the process. H2C03° + M£2+ = MgHC03+ + H+ (8) 136 Figure 28, Variation of hydrogen-containing species with, temperature in the Atlantis XX brine. L O G a | -2 -3 -4 -5 HS" 250 100 150 200 50 O TEMPERATURE (°C) 138 The distribution of dissolved sulfur-bearing species (Fig. 29) shows that NaSO^~ is of overwhelming importance except at very high temperature and free sulfate ion is quite low as a result. An important factor in this model (Fig. 29) is the variation of a^ g with temperature for the fD values used. H~S remains very low compared to total 2 dissolved sulfate throughout and actually increases slightly as temperature falls. H^S concentration at 60° C would be approximately 2.0 ppm, which would probably be too low to detect. Also, Figure 29 shows that sulfide species disappear completely due to oxidation upon cooling to 25° C. Thus, there would be little hope of observing dissolved sulfide in a water sample which has undergone the pressure and tempera ture changes inherent in raising it through 2,000 m of water. However, the present model does explain the blacken*)* ing of a new brass messenger observed by Miller et al. (1966) after hydrocasts into the hot brine deep* **2^ present but in low quantities and is not likely to be detected except by in situ sampling and analysis. More importantly, consideration of free metal ions and metalchloride complexes (Fig. 30) indicates that 2+ 2+ + 2 + Zn , Pb , Cu (not shown), and Fe activities increase with falling temperature. In general, copper, zinc, and lead chloride complexes show slight decreases in activity with temperature. This is due to activity coefficient ________________________________________________________________139 Figure 29m Variation of sulfur speciation with, temperature in the Atlantis XX brine# L O G c h -2 -4 -5 HS" 100 150 200 250 50 O TEMPERATURE (°C) 141 Figure 30* Variation of chloride complexes with temperature in the Atlantis XX brine. L O G Q i -2 -3 2* -4 ZnCI* -5 ZnCI CuClj -6 ZnCI PbC i; CuCI1 PbCI® -71 PbCI -8 250 -9 -10 200 50 xhj effects and, in fact, molalities of chloride complexes increase with temperature. Since chloride complexes make up the overwhelming percentage of total copper, lead and zinc, this small change in molality of complexes is reflected by a large change in free-metal ion activity (Fig. 30). Ferrous iron dominates the iron species over the 2 + Conditions of this model. However, Fe activity also decreases with increased temperature, due mainly to the * ! » stability of FeOH at higher temperatures. Mineral Solubilities In order to evaluate transport and precipitation processes in the Red Sea geothermal system, it is neces sary to consider possible equilibria with solid phases. Solubility relationships of some important solid phases are plotted (Fig. 31) sis a function of temperature and the logarithm of the ionic activity product (lAP) divided by the solubility product (Ksp). A positive value of log (lAP/Ksp) thus indicates that the solution is over saturated at the given temperature, while negative values indicate undersaturation. The solubility relationships for sulfide minerals (Fig. 31) show that above 150° C the brine becomes under saturated and thus could act as a transport system, solubilizing metals during subsurface interaction with 144 Figure 31* Solubility of important hydrothermal minerals in the Atlantis XX brine. 145 L O G (IAP/K Galena Chalcopyrite Anhydrite FeS Barite -2 Sphalerite Siderite -3 -4 UNDERSATURATED -5 -6 -7 200 100 150 250 50 O TEMPERATURE (°C) 146 rocks. Upon cooling the brine becomes oversaturated first with respect to chalcopyrite, then galena, sphalerite, and iron monosulfide. This model for sulfide transport and deposition along temperature gradients is thus in very- good agreement with observed processes in the Atlantis XI Deep. One puzzling feature, however, is the relative lack of galena in the hydrothermal sediments. Pyrite is persistently and very greatly supersaturated (Table XII) throughout the whole range of temperatures. It is interesting to note that simple cooling of the brine to 25° C (without changing fn ) results in chalcopyrite, 2 sphalerite, and galena becoming undersaturated. Pyrite remains greatly supersaturated. This solubility calcula tion, thus, is in excellent agreement with the observation the sphalerite and possibly chalcopyrite disappear at the expense of pyrite during core storage. Also, petrographic observation of the pyrite indicated it to be quite coarse grained compared to the other sulfide minerals. Again this is suggestive of recrystallization to form the pyrite. It is difficult to evaluate if pyrite ever occurs as a primary precipitate in the hydrothermal sediments. Bischoff (1969) found very little pyrite in the fresh sediments he examined. Without further detailed mineral- ogical information, on fresh and carefully handled samples, it must be assumed that pyrite is a very minor 147 Table XII, Degree of saturation (log IAP/Ksp) of the hot brines Mineral ATLANTIS II DEEP, 60° C Brine 250° ATLANTIS II DEEP 50° C Brine Layer 25° 60° 100° 150° 200° 25° 50° BaSO. 4 -0,05 -0,74 -1.14 -1,41 -1,48 -1,60 FeC03 -0,93 -0,74 -1,15 -1,94 -2,38 -2,46 -3,06 -2,74 CaC03 -1,28 -1,24 -1.74 -2,58 -3,09 -3,09 -1,14 -0,91 CaSO. 4 -0,61 -0.55 -0,27 +0.07 40.44 40.71 -0,42 -0,36 CaMg(C03)2 -2,71 -2,47 -3,39 -5,04 -6,06 -6,17 -1,94 -1,37 CuFeS2 -18,95 +8,44 +3,77 -0,85 -4,64 -8,02 -12,8 +6,59 PbS -12,08 +2,43 -0,30 -2,84 -4,84 -6,22' -7,69 +2,50 ZnS +12,32 +2,31 -0,43 -3,11 -5,76 -7.47' -7,66 +2,71 FeS -15,44 +0,83 -0,26 -1,35 -2,14 -2,61 -13,15 -1,63 FeS2 +131,25 +146,6 +133,54 +121,2 +111,9 +104,1 +136,5 +147,1 148 component as a primary precipitate from the brine. Perhaps primary pyrite precipitation is prohibited by strong kinetic barriers as in recent marine sediments (Berner, 1970). Other important hydrothermal phases plotted in Figure 31 are barite, anhydrite, and siderite. Barite and siderite are close to saturation at lower temperatures and in fact barite plots exactly at saturation at 25° C, The hydrothermal carbonate actually found in the Atlantis XX sediments is a manganosiderite solid solution for which thermodynamic data are not available. Anhydrite is close to saturation throughout, becoming slightly supersaturated at higher temperature, which implies that any primary anhydrite in the deposit should be deposited very close to the brine vent. Sulfur Isotopes and Species Distribution The sulfur isotope ratio of each dissolved sulfur specie can be calculated as outlined in Ohmoto (1972) using theoretically derived fractionation factors (Fig, 32), Fractionation factors between sulfate complexes are assumed to be negligible. The sulfur isotope ratio of source material (6 S _ ) is assumed to be +20 °/oo, for v Z s' the presumed source of contemporary sea water sulfate. This assumption would not be substantially changed if the source material is Miocene evaporite sulfate minerals. 149 Figure 32. Sulfur isotope fractionation factors (after Ohmoto, 1972). (°C) 600 400 200 100 50 25 + 30 80 f 25 + 20 60 40 20 + 10 + 3 ro C O co HS [PbSJ 10 15 12 151 Actual calculations were carried out using the fol lowing methods• Distribution of* sulfur isotopes among sulfur species is described by: £S=^SH2SXH2S+5SHS_+$SS2-XS2-+$S 0 XCS04 (8) where 2 Xi=m./ (9) X. is a conventional mole fraction term, ZTS is total 1 ’ sulfur, and XT SO^ is total sulfate species* Xsotopic enrichment factors with respect to H^S are derived from Figure 32 and are expressed as follows: 6Si = g S H 2 S + A ± ( 1 0 ) Where /S. ^ is the per mil difference between species i and S c. Equations (8) and (lO) can be combined and 2 solved for & S„ 0 as: 2 j-o ZS-( A HS_XHS-+i^S-XS-+Aso,,2 X £S0.) (ll) asH S ------------------------------------------— * h 2 s + x h s " + x s 2 + x -zsok ^ and Ai's are known and X. can be easily A - b 1 calculated from the species distribution. Once & S,. 0 is 2 calculated, £ values for other species can be determined by back substitution in equation 10. Sulfur isotope ratios calculated in this manner (Fig. 33) indicate that isotopic exchange equilibrium is not adequate to explain the +5 °/oo average ratios of 132 sulfide materials in the Atlantis XX Deep, Equilibration temperatures of* 350°-400° C would be necessary for the observed bS ratios# Xf sulfur isotope equilibration occurred in the incoming brine, even at temperatures as high as 200° C, the expected sulfur isotope ratio of Q / precipitated sulfides would be about -15 /oo, Nonetheless, the fact that sulfate in the brine has a sulfur isotope ratio of unaltered sea water (+20 °/oo) requires that only a very small fraction of the total sulfate be reduced to sulfide# Otherwise the residual 34 sulfate would be greatly enriched in S • Since it is unlikely that the brine was ever at 350° C, sulfides must form by a process where isotopic exchange equilibrium is not attained# 153 DISCUSSION AND CONCLUSIONS Origin of Atlantis II Sulfides The geochemical model discussed above suggests that conditions very similar to those presently observed in the brine pool may be adequate to account for sulfide trans port and deposition* Before inferences can be made concerning facies alternation and sulfide deposition in similar ancient deposits, it is necessary to quantita tively evaluate present processes* An estimate of how much sulfide is being delivered to the Atlantis II Deep by the present brine can be derived if brine flux and sulfide concentration are known* Schoell (1975) calculates an approximate rate of brine influx of 2.4 x 10^ liters/minute. This estimate is based on observed temperature increase and possible volume increase of the brine pool since 1964* If the incoming brine is significantly hotter than 210° C, then the flux estimate would be accordingly lower. The Na-K-Ca geo thermometry, however, indicates that the brine is not significantly above 200° C* Also, even if the temperature of the incoming brine was as high as 300° C, the flux estimate would be of the same order. 154 An estimate of* the rate of* sulfide precipitation can be derived assuming that the incoming brine is at 150-200° C and that all the sulfide brought in is precipi tated, This seems a reasonable assumption since the 60° brine is greatly oversaturated with respect to the sulfide —4 2 minerals. The concentration of H^S is about of 10 # M i q o and flux of 1.6 x 10 Kg brine/10 years would precipitate 11 2 9#6 x 10 g sphalerite/10 years. This amount of sphalerite would constitute about 0.3 percent by weight of the total sediment being deposited, assuming a sedi- O mentation rate of 75 cm/10 years, 1.5g/cc bulk density, and a depositional area of 42Km (present brine pool)• The calculated percentage of sphalerite in the sedi ment is somewhat lower than that observed in the iron— montmorillonite facies by Bischoff (1969) which ranged up to a few percent. If cooling is the cause of deposition, however, then it is likely that sulfide facies deposition be localized near the brine vent. In fact, sulfides are probably precipitated in the ascending brine and are brought into the deep as particulate matter. This factor may explain large variations in thickness and extent of the sulfide facies from basin to basin, as observed in Figure 16. Xt may be that while sulfides are precipitated rapidly near the vent, more oxidized material, which settles out uniformly from the transitional brine layer, predominates in other areas. Recalculation of sphalerite _______________________________________________________________ 155_ 2 content, assuming a lOKm depositional area, gives 6 per cent by weight in the sediment# Thus, it appears that this model can successfully account for precipitation of sulfides in the iron—montmorillonite facies# r Deposition of the sulfide facies, which contain up to 40 percent sulfides by weight, could have resulted from a number of different processes# The simplest and most attractive explanation is that sulfide minerals greatly predominate in areas of very close proximity to the vent# Thus, deposition of massive sulfide beds is very localized due to differential precipitation and sedimentation of sulfide minerals# A second possible explanation might be increased sulfide production in the brine due to lower fn conditions 2 than at present# This could occur as a result of brine interaction with organic-containing shales as suggested by Kaplan ejfc al# (1969)# However, if a large percentage of the sulfate were reduced, then the residual sulfate would 3 4 be quite enriched in S # As mentioned earlier, iso- topically heavy sulfate minerals which are clearly a result of brine precipitation have not been observed# — 3 5 Nonetheless, an increase in sulfide to about 10 * molal S 34 would not significantly alter the aS ratios of the remaining sulfate and could account for deposition of the sulfide facies if limited to a relatively small (r - * lOKm ) area# 156 Another alternative process for precipitating massive sulfide beds is simply an increased brine flux. This would require an increase of about an order of magnitude over present influx and it is difficult to evaluate if this has ever occurred. As discussed in an earlier section, how ever, there is stratigraphic evidence that both the SU^ and SXJ^ zones were deposited when the brine pool was at its maximum extent. At present it is impossible to evaluate which of these mechanisms was responsible for the formation of relatively pure sulfide zones and, in fact, all three factors may be important. Xt seems likely, though, that differential settling is the controlling process. This would account for the difficulties in correlating age- dated sulfide zones and would not require radical changes in brine chemistry or plumbing of the system. Sulfur Isotope Ratios of Sulfides As mentioned above, the apparent fractionation factor of 15 °/00 between brine sulfate and precipitated sulfide minerals is difficult to explain by isotopic exchange equilibria. The theories of sulfide formation proposed by earlier workers (Kaplan e^t al. , 1969; Craig, 1969; Watson and Waterbury, 1969) all required special conditions unique to the Red Sea system. There is, how ever, abundant evidence that this deposit is isotopically ________________________________________________________ 157 analogous to many ancient massive sulfides and in spite of all the possible variables which effect sulfur isotope ratios (Rye and Ohmoto, 197*0 * ^^rie similar isotopic compositions imply a common and simple mode of origin for this type of deposit. Holser and Kaplan (1966), Thode and Monster (1965)* and Nielsen (1965) have shown that sea water sulfate, as preserved in sulfate-bearing evaporite deposits, has varied remarkably in z> S ratio throughout geologic time. Sangster (1968, 1971) has further noted a striking paral- lelism between <&S ratios m evaporites and contempora neous massive sulfide deposits. His data compilation is divided into volcanic and sedimentary types on the basis of host rock lithology (Fig. 33)• Sulfur isotope data from 100 massive sulfide deposits were compiled to make these plots and the evidence of a genetic connection between sea water or evaporites and this type of sulfide deposit is somewhat compelling. Sangster suggests that, in general, such massive sulfides are a result of bacterial sulfate reduction. The observed average sulfate-sulfide fractionation factors (A ^ S34) are 17«3 °/00 for volcanic types and 13*6 °/oo for sedimentary types. He further proposes that the larger A ^ S34 for deposits of volcanic association is due to admixture of volcanic sulfur (0 °/00) into the sulfides of the deposits. Sasaki (1970), however, points 158 Figure 33* Sulfur isotope ratios of sulfur species in the Red Sea brine. Distribution of dissolved sulfur species from Figure 29 and fractionation factors from Figure 32. ♦20 + 1 0 - 1 0 - 2 0 - 5 0 - 6 0 200 1 5 0 1 0 0 TEMPERATURE °C TOTAL 160 out the Improbability of this suggestion. Volcanic sulfur has probably remained close to 0 °/oo throughout geologic time while "volcanic type" massive sulfide averages have varied from +15 °/oo (Devonian) to -10 °/oo (Mississip- pian). It is clear that addition of volcanic sulfur would A 3k not result xn a systematic variation of /\d>S . In addition, sulfur isotope studies of recent sedi ments (Kaplan et al., 1963; Goldhaber and Kaplan, 197^> and Hartmann and Nielsen, 1969) have indicated that bacterial fractionation factors are much larger than 15 °/oo. It is possible that bacteria have produced these relatively small isotopic fractionations, as they ap parently do in petroleum (Thode and Monster, 1 9 6 5)> but it seems more likely that hydrothermal processes are involved, especially in deposits of volcanic association. Sasaki (1970) has suggested that Sangster*s (1968, 197l) observations are a result of isotopic exchange equilibria. Much recent work has been devoted to inter pretation of hydrothermal sulfides using such equilibria (Rye and Ohmoto, 197^0* Isotopic fractionation factors (Sakai, 19685 Ohmoto, 1972) have been derived for most important aqueous species and minerals by theoretical and experimental methods (Fig. 16). As emphasized by Ohmoto (1972), sulfur isotope ratios are a function not only of temperature but also of the relative concentration of dissolved species, which can 161 Figure 3^0 Variation of average sulfur isotope ratios in massive sulfide ore deposits with geologic time (Sangster, 1971) . 162 tsMX, +30 +20 +10 0 -10 i 1 i ----i —! --1 A(SS ) L. CRETACEOUS * 9-4 PERMIAN V 11-3 PENNSYLVANIAN ^ 20 0 MISSISSIPPIAN > X 11.0 U.DEYONIAN ^ 10. 2 SILURIAN ^ Y 15<5 ORDOVICIAN d X 19-5 J-----------» 13*8: 2 1 -5 0 2 1 -7 5 2 2 0 0 22-25 2 2-50 S3 /S34 AVGE. 163 be described using various chemical parameters (fn , pH, 2 fc )# It should be pointed out, however, that use of 2 isotopic exchange relationships requires equilibrium processes in a closed system. The closed system require ment implies that the amount of sulfur deposited must be insignificant compared to the total sulfur reservoir in the system. Comparison of Sangster’s generalized relationships to known fractionation factors (Fig. 32) suggests that geo logically unreasonable temperatures (400°-500°) would be required to account for the observed sulfate-sulfide fractionations. In addition, the sedimentary deposits (A S S"^ ss 13.6 °/oo) should have formed at higher tempera ture than the volcanic deposits (A 8 = 17*3 °/oo). Using sulfur isotope systematics as described by Ohmoto (1972), one can derive the desired isotope ratios at lower temperatures simply by adjusting pH and fn . For 2 o 3^ + o / example, if o S of sea water is +20 /00 and thxs represents total sulfur, total ionic strength is 1.0 m, o tr 3h o / and temperature is 250 C, then pyrite with O S = 5 /oo can be precipitated if fn = 10 and pH = 5*5 (Ohmoto, 2 1972). However, it is unreasonable to assume that in IOO ore deposits from widely varying environments, conditions (T, f0 , pH, r S) were often adjusted to give a 15 °/oo 2 difference between total sulfur and sulfides. Moreover, 164 in the situation described above, residual S0^= would have _ q / | , q , a S S ratio of about +30 /oo. An examination of co existing sulfate minerals in massive sulfide deposits where paleotemperatures are well known provides sub stantive evidence of the actual fractionation processes. The Kuroko ores of Japan are outstanding examples of relatively unaltered massive sulfides of volcanic as sociation (Tatsumi, 1965* Lambert and Sato, 197*0# These deposits are Miocene age and Kajiwara (1973a* h) has sug gested that contemporaneous sea water was the ore-bearing fluid. Sulfur isotope studies (Kajiwara, 1971* Kajiwara and Date, 197l) show that 6 S^^ values of sulfides generally range from +2.4 to +8,2 °/oo. Kajiwara (1971) has interpreted these data in terms of fQ -pH conditions, 2 as in Ohmoto (1972). He concludes that the sulfides were precipitated by an isotopic equilibrium process at 2J0° C with sea water providing the sulfur source. Barite and gypsum coexisting with Kuroko sulfides range from +21,5 to +23,4 °/oo. As discussed by Kajiwara (l97l)* sulfates precipitating from the same ore solution should have S ratios of at least +28 °/oo. The observed +22 °/oo ratios thus indicate isotopic exchange disequilibrium. Kajiwara suggests later interstitial precipitation of barite but this seems tenuous, Sakai et al. (1970) have confirmed that Kuroko-associated sulfates typically fall in a range of +21 to +24 °/oo. __________________________________________________________ 165 Other well documented examples of massive sulfide deposits of this type where sulfur isotopes have been measured in coexisting sulfates and sulfides, and deposi- tional temperatures are well defined, are scarce in the literature. Some natural examples of apparent dis equilibrium are discussed, however, by Ohmoto (1972) and Bye and Ohmoto (197*0# Other examples of hypogene ore deposits where sulfate-sulfide disequilibrium may have occurred are given in Sakai (l957)> Jensen (1959), Ault and Kulp (i960), and Gavelin ejb al. (i960). Experiments in the laboratory (Puchelt, 1969* and Voge and Libby, 1937) suggest that the rates of isotopic exchange reactions, particularly between reduced and oxidized sulfur species, are extremely slow below 200° C. These results were confirmed recently by Robinson (1973)# He found that sulfur hydrolysis equilibria, rapid at 230° C, could only be approached from one direction at 200° C. Striking confirmation of isotopic disequilibrium is provided by detailed studies of sulfur isotopes in the natural geothermal systems of New Zealand. Rafter and Wilson (1958) analyzed sulfur isotopes in waters from Wairakei. Average sulfur isotope ratios of dissolved sulfate and sulfide are +21.4 °/oo and +4.4 °/oo, re spectively. Maximum in situ temperature is 2J0° C, much lower than expected for the observed +17*0 °/oo dif 166 ference. Comparison to Figure 32 indicates an equilibra tion temperature of 370° C. Steiner and Rafter (1966) have confirmed that pyrite and pyrrhotite from core samples at Wairakei are isotopically similar to the dissolved H^S measured by Wilson# In summary, it appears that sulfate-sulfide isotopic exchange is prohibitively slow in a number of natural and experimental hydrothermal systems. This observation has potentially important implications to the study of ancient sulfide metallogenesis. The methods outlined by Ohmoto (1972) have been used repeatedly in the last few years (see, for example, Rye and Ohmoto, 197*0 "t 0 predict the isotopic composition of the sulfur source. This method depends on the calculated ratio of reduced to oxidized sulfur species in the ore fluid and the established fractionation curves (Fig. 32). If exchange equilibria is incomplete, then suggested sulfur sources for low- temperature hydrothermal deposits may be incorrect. Laboratory study of bacterial sulfate reduction at room temperature provides some interesting insights into possible sulfide-forming processes in the Red Sea geo thermal system. Jones and Starkey (1957) and Kaplan et al. (i960) have shown that sulfate-reducing bacteria produce H^S depleted in S , but rarely by more than 15 /00 under normal growth conditions at 25° C. More recently, how ever, Kaplan and Rittenberg (1964) have measured fraction- _______________________________________________________________ 167 ations as large as 46 °/oo under different metabolic conditions# The exact processes involved in sulfate reduction in these two sets of experiments is still controversial but it is clear that a kinetic isotope ef fect is involved rather than exchange equilibria# The latter would require sulfate-sulfide fractionation of 73 °/oo. Harrison and Tbode (1958) have suggested that If the rate limiting step in sulfate reduction is the initial cleavage of the S-0 bond to form sulfite, then the maximum theoretical fractionation factor is 25 °/oo due to the kinetic isotope effect# Inorganic sulfate reduction experiments at room temperature (Harrison and Thode, 1957) agreed with this conclusion# Unfortunately, kinetic isotope effects have not been measured at higher tempera tures than 25° C and isotopic exchange equilibria have not been substantiated at temperatures below 200° C. Thus, the only reasonable conclusion is that in the Red Sea, in many other ancient sulfide deposits, and in active geothermal systems, the sulfur isotope ratio of reduced sulfur species is controlled not by kinetic isotope effects which yield an apparent fractionation factor of 10-20 °/oo# It appears that, in attacking the problem of massive sulfide formation, Sangster (1968, 1971) chose a group of deposits which form in the temperature range of 50-300° C. This is precisely the temperature 168 range where isotopic exchange equilibria is doubtful and kinetic isotope effects need to be seriously evaluated. Xn summary, the Red Sea sulfides can be accounted for by the following model. The brine source is normal Red Sea water which interacts with evaporites at very low temperatures to attain high salinity (no oxygen isotope shift expected at low temperature). The brine is then circulated and heated to about 200° C. Oxygen isotope shift is precluded by high water-rock ratios. The driving force for circulation is probably heated recent intrusives in the rift zone and fn in the brine is controlled by 2 magnetite—hematite equilibria. A small amount of dis solved sulfide forms through sulfate reduction with sulfur isotope ratios being controlled by kinetic isotope ef fect. Sulfides are precipitated as the brine cools in the Atlantis XX Deep. As previously described, the base-metal sulfides in the Atlantis II Deep show variations in the range of about +1 to +15 °/oo. Most of the values, however, center around +5 °/oo. In certain hydrothermal ore deposits, Rye and Ohmoto (197*0 have attributed such sulfur isotope variations to fluctuations of fn and pH at deposition. 2 Such conclusions are justified where there is evidence of isotopic equilibrium or where there are overall trends within the deposit. In the Red Sea, however, variations in the sulfide facies are not systematic. Also, the 169 calculated sulfate/sulfide ratio of about 10^ indicates that sulfur isotope ratios would be insensitive to minor changes in fn even if isotopic equilibria were attained. 2 Alternatively, sulfur isotope variations in sulfides from the Red Sea system may be due to admixture of bacterial sulfides produced at the brine-sea water inter face. However, even in sediments from the iron- montmorillonite facies where sulfides constitute less than 1 percent of the sediment, sulfur isotope ratios of +4 °/oo are observed (Kaplan jet al., 19^9 5 Hartmann and Nielsen, 1966). It therefore is very unlikely that any significant amounts of bacterial sulfides are falling in from the water column. The most reasonable explanation of the observed ^ 34 fluctuations in b S ratios is variation in the degree of isotopic exchange. Variations such as these, where various degrees of equilibria are attained, would be expected for the proposed non-equilibrium process. In c 34 fact, it is somewhat surprising that the O S ratios are as constant as observed but this seems to be characteristic of this type of deposit. Origin of Atlantis II Sulfates Anhydrite in the Atlantis II Deep occurs predomi nantly in two modes: as crystalline infillings of voids and fractures within the sediments and as discontinuous ________________170 nodules, lumps, and beds, A limited number of* sulfur isotope determinations on tbe latter have indicated that 3** they are slightly enriched in S and may be related to isotopically similar (Shanks et al,, 197*0 Miocene evaporites which are believed to outcrop in the Atlantis IX Deep, Crystalline aggregates of anhydrite, which occur as Jfissure injfillings in the Southwest Basin, are believed to be related to active brine venting. The geochemical model calculations indicate that anhydrite should precipi tate at brine temperatures greater than about 120° C and thus agrees with the geologic observations. At lower temperatures even in the 50° brine layer (Table XI) which contains considerably more sulfate (Table l) anhydrite is undersaturated. Thus, anhydrite cannot precipitate by mixing of the brine and Red Sea deep water, as suggested by Bischoff (1969)9 and massive anhydrite in the Deep most likely is mechanically introduced. The discovery of ubiquitous barite as a minor component in the sulfide facies provides important informa tion about brine conditions at deposition. Sulfur isotope ratios of barites are heavier than coexisting sulfides by erratically varying amounts but are usually lighter than brine sulfate or primary anhydrite. These ratios suggest that the barite is forming in situ by oxidation of sulfides, either as particulate sulfides settle out or ______________ 171 after deposition. If barite was precipitating directly from the brine, its sulfur isotope ratios would be invari ably +20 °/oo. Earlier considerations of solubility relationships, however, indicated that the brine would be saturated with barite at 25° C. Thus, the possibility exists that the barite has formed during core storage. In this study, barite sulfur ranged between 1,0 and 0.1 percent in the sediment. If the barite has formed during core storage, the only source of barite would be the interstitial water. Barite concentration in the brine is only 0.9 mg/Kg and thus it is unreasonable that even the smallest amount formed after sample collection. The conclusion that sulfides are, in small amounts, altering to barite has two important consequences. First, it provides compelling support for the assumptions that the brine is in sphalerite-barite equilibrium, as in the fQ calculations. The second important point is that 2 barite is forming from a supposedly undersaturated brine, indicating that there may be a problem with the analytical 2 + data for Ba or with the calculated free S0^= activity. Summary As a result of this and other recent studies, the processes of ore formation in the Red Sea hydrothermal system can now be assessed in considerable detail. The _____________________________________________ 172 most important features can be summarized as follows: 1 g 1. 5 D and £ 0 relationships indicate that the Atlantis XI brine is a result of circulating Red Sea waters. The exact source of the waters is not clear. Lack of isotopic exchange indicates a large water/rock ratio during subsurface circulation. 2. Independent chemical and physical evidence shows that brine temperature is at least 200° C as it flows into the Atlantis XI Deep. 3# Hydrothermal activity has occurred at other locations along the axial rift zone of the Red Sea at various times in the past, notably in the Thetis Deep. 4. Massive sulfide beds were deposited in the Atlantis II Deep as early as 20,000 years ago. 14 C ages show that lithostratigraphic units are not synchronous throughout the Atlantis II Deep, suggest ing lateral facies differentiation during precipitation. 6. Overflows have occurred at various times, contributing hydrothermal sediments and brine to the neighboring Albatross, Shagara, Wando, Discovery, and Chain deeps. 7• Base-metal sulfides are deposited from brine which is compositionally similar to the present brine. Precipitation is caused by cooling. 8. Equilibrium sulfate-sulfide sulfur isotope fractionation factors are not obtained in the Red Sea or _______________________________________ 173 in many other low-temperature hydrothermal systems# Rather an apparent fractionation of 15 °/oo due to the kinetic isotope effect is observed# 174 REFERENCES 175 REFERENCES Ault, W. U. and J. L. Kulp, i960, Sulfur isotopes and ore deposits: Econ, Geol., v. 55> p. 73-100. Anderson, C. A., 1969* Massive sulfide deposits and volcanism: Econ. Geol., v. 64, p. 129-146. Bachinski, D. J., 1969* Bond strength and sulfur isotopic fractionation in coexisting sulfides: Econ. Geol., v. 64, p. 56-65# • • Backer, V. H. and M. Schoell, 1972, New deeps with brines and metalliferous sediments in the Red Sea: Nature, Physical Science, v. 240, p. 153-158. Backer, V. H. and H. Richter, 1973> Die rezente hydro- thermalsedimentaire lagerstatte Atlantis XI - Tief im Roten Meer: Geol. Rundschau, v. 62, p. 697-741. Baumann, A., H. Richter, and M. Schoell, 1973> Suakin Deep: brines and hydrothermal sediments in the deep est part of the Red Sea: Geol. Rundschau, v. 62, p. 684-697. Berner, R. A., 1970, Sedimentary pyrite formation: Amer. Jour. Sci., v. 268, p. 1-23# Bignell, R. D., 19759 Timing, distribution and origin of submarine mineralization in the Red Sea: Trans. Inst. Min. Metall., v. 84, p. B1-B6. Bischoff, J. L., 1969> Red Sea geothermal brine deposits: their mineralogy, chemistry, and genesis: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer-Verlag New York, Inc., p. 368-401. ________, 1972, A ferroan nontronite from the Red Sea geo thermal system: Clays and Clay Minerals, v. 20, p. 217-233. 176 Bischoff, J. L. and F. T# Manheim, 1969, Economic potential of* Red Sea heavy metal deposits: in Degens, E. T. and D. A. Ross (eds,), Hot brines and recent heavy metal deposits in the Red Sea, Springer-Verlag New York, Inc,, P. 535-541. Bischoff, J. L. and F. ¥. Dickson, 1975* Seavater-basalt at 200° C and 500 bars: Implications for origin of* sea-floor heavy-metal deposits and regulation of sea water chemistry: Earth Planet. Sci. Letters, v* 25, p. 385-397• Blissenbach, E. and R. Fellerer, 1973» Continental drift and the origin of certain mineral deposits: Geol. Rundschau, v. 62, p. 812-840. Bostrom, K., M. N. A. Peterson, 0. Joensun, and D. E. Fisher, 1969* Aluminum-poor ferromangoan sediments on active oceanic ridges: Jour. Geophys. Res., v. 74, no. 12, p. 3261-3278. Bostrom, K., 0. Joensun, S. Valdes, and M. Riera, 1972, Geochemical history of south Atlantic Ocean sediments since late Cretaceous: Mar. Geol., v. 12, p. 85-121. Brewer, P. G., J. P. Riley, and F. Culkin, 19^5, The chemical composition of the hot salty water from the bottom of the Red Sea: Deep-Sea Res., v. 12, p. 497-503. Brewer, P. G. and D. ¥. Spencer, 19^9* A note on the chemical composition of the Red Sea brines: in Degens, E. T. and D. A. Ross (eds.). Hot brines and recent heavy metal deposits in the Red Sea, Springer- Verlag New York, Inc., p. 174-179. Brewer, P. G., C. D. Densmore, R. Munns, and R. J. Stanley, 1969, Hydrography of the Red Sea brines: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer- Verlag New York, Inc., p. 138-147. Brewer, P. G., T. R. S. ¥ilson, J. ¥. Murray, R. G. Munns, and C. D. Densmore, 1971» Hydrographic observations on the Red Sea brines indicate a marked increase in temperature: Nature, v. 231, p. 37-38. Bruneau, L., N. G. Jerlov, and F. F. Koczy, 1953* Physical and chemical methods: in H. Petterson (ed.), Reports of the Swedish deep-sea expedition, v. 3» p. 101-112. 177 Clayton, R. N., 1959* Oxygen isotope fractionation in the system calcium carbonate-water: Jour. Chem. Phys., v. 30, p. 1246-1250. Clayton, R. N., L. J. P. Muffler, and D. E. White, 1968, Oxygen isotope study of calcite and silicates of the River Ranch, No. 1 well, Salton geothermal field, California: Amer. Jour. Sci., v. 266, p. 968-979# Clayton, R. N. and A. Steiner, 1975, Oxygen isotope studies of the geothermal systems at Wairakei, New Zealand: Geochim. Cosmochim. Acta, v. 39, P# 1179- 1186. Constantinou, G. and G. J. S. Govett, 1973, Geology, geochemistry, and genesis of Cyprus sulfide deposits: Econ. Geol., v. 68, p. 843-858. Craig, H., 1963, The isotopic geochemistry of water and carbon in geothermal areas: from Conference on Nuclear Geology in Geothermal Areas, Spoleto, Italy, ed. Tongiorgi, p. 17-53. ________, 1966, Xsotopic composition and origin of the Red Sea and Salton Sea geothermal brines: Science, v. 154, p. 1544-1548. ________, 1969, Geochemistry and origin of Red Sea brines: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer- Verlag New York, Inc. Cronan, D. S., T. G. van Andel, G. R. Heath, M, G. Dinkelman, R. H. Bennett, D. Bukry, S. Charleston, A. Kaneps, K. S. Rudolfo, and R. S. Yeats, 1972, Xron-rich basal sediments from the eastern equatorial Pacific: Leg 16, Deep Sea Drilling Project: Science, v. 175, P# 61-63. Degens, E. T. and D. A. Ross, 1969 (eds.). Hot brines and recent heavy metal deposits in the Red Sea, Springer- Verlag New York, Inc. . 1970, The Red Sea hot brines: Sci. Amer., v. 222, no. 4, p. 32-42. Dietrich, G. and G. Krause, 1969, The observations of the vertical structure of hot salty water by R. V. Meteor: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer- Verlag New York, Inc., p. 10-14. 178 Drake, C. L. and R. ¥. Girdler, 1964, A geophysical study of* the Red Seas Geophys. Jour, Roy, Astron, Soc, , v. 8, p, 473-495, Duke, N, A. and R. ¥. Hutchinson, 197^, Geological relationships between massive sulfide bodies and ophiolitic volcanic rocks near York Harbour, Newfound lands Canadian Jour, Earth Sci,, v, 11, p, 53-69. Faber, E. and M, Schoell, 1975, Oxygen and hydrogen isotopic composition of Red Sea brines. In prepara tion, Fisher, F, H, and A, P, Fox, 1975, NaSO^” ion pairs in aqueous solutions at pressures up to 2000 atmospheress Jour, Solution Chem,, v, 4, p, 225-236. Fournier, R. 0. and A. H. Truesdell, 1973, An empirical Na-K-Ca geothermometer for natural waterss Geochim. et Cosmochim. Acta, v. 37, P. 1255—1275. Friedman, G. M., 1972, Significance of Red Sea in problems of evaporites and basinal limestoness Amer. Assoc. Petr. Geol. Bull., v, 53, p. 870-883. Friedman, I. and K. Hardcastle, 1974, Deuterium in inter stitial waters from Red Sea cores: Initial Reports, Deep Sea Drilling Project, v. 23, p. 969-970. Garrels, R. M. and M. E. Thompson, 1962, A chemical model for seawater at 25° C and one atmosphere total pressure: Amer. Jour. Sci., v. 260, p. 57-66. Gavelin, S., A. Parwel, and R. Ryhage, i960, Sulfur isotope fractionation in sulfide mineralization: Econ. Geol., v. 55, P. 510-530. Gevirtz, J. L. and G. M. Friedman, 1966, Deep-sea carbonate sediments of the Red Sea and their implications on marine lithification: Jour. Sed. Petrol., v. 36, p . 143 - 152 . Giggenbach, ¥. F., 1971, Isotopic composition of waters of the Broadlands geothermal field: New Zealand Jour. Sci., v. 14, p. 959-970. Goldhaber, M. B. and X. R. Kaplan, 1974, The sedimentary sulfur cycle: The Sea, Interscience, v. 5. Hackett, J. P., 1972, Stratigraphy, extent, and geologic history of the Red Sea geothermal deposits: Un- _________________________ 179 published Master of Science thesis, University of Southern California* Hackett, J* P. and J* L. Bischoff, 1973* New data on the stratigraphy, extent, and geologic history of the Red Sea geothermal deposits: Econ* Geol*, v* 68, p* 533- 564. Harrison, A* G* and H* G. Thode, 1957* The kinetic isotope effect in the chemical reduction of sulfate: Trans. Faraday Soc., v* 53* P# 1648-1651# * 1958, Mechanism of bacterial reduction of sulfate from isotope fractionation studies: Trans. Faraday Soc*, v* 54, p. 84-92. Hartmann, M. and H. Nielsen, 1966, Sulfur isotopes in the hot brine and sediment of Atlantis XX Deep (Red Sea): Marine Geol., v. 4, p. 305-306. ________, 1969, S34 werte im rezente meeressedimenten und ihre deutung am beispiel eineger sedimentprofile aus der westlichen ostsee: Geol. Rundschau, v. 58, p. 621-655# Hartmann, M., 1970, Bericht uber geochemische unter- suchungen in den hydrothermallangenbechen am Boden des Roten Meeres: Geol* Rundschau, v. 60, p* 244-2 56* . 1973* Untersuchung von suspendiertem material in den hydrothermallagen des Atlantis II tiefs: Geol. Rundschau, v. 62, p. 742-754. Helgeson, H. C., 1964, Complexing and hydrothermal ore deposition: Pergamon Press, New York-London, 128 p. ________, 1969, Thermodynamics of hydrothermal systems at elevated temperatures and pressures: Amer. Jour. Sci., v. 267, P# 729-804. Hendricks, R. L., F. B. Reisbick, E. J. Mahaffey, D. B. Roberts, and M. N. A. Peterson, 1969, Chemical composition of sediments and interstitial brines from the Atlantis II, Discovery, and Chain deeps: in Degens, E. T. and D. A* Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer- Verlag New York, Inc., p* 407-440. 180 Holland, H. D., 1965, Some applications of thermochemical data to problems of ore deposits. XX. Mineral as semblages and the composition of ore-forming fluids: Econ. Geol., v. 60, p. 1101-1166. Holser, ¥. T. and I. R. Kaplan, 1966, Isotope Geochemistry of sedimentary sulfates: Chem. Geol., v. 1, p. 93- 135. Hutchinson, R. ¥. and D. L. Searle, 1971* Stratabound pyrite deposits in Cyprus and relations to other sulphide ores: Soc. Mining Geol. Japan, spec, issue 3, p. 198-203. Hutchinson, R. ¥., 1973* Volcanogenic sulfide deposits and their metallogenic significance: Econ. Geol., v. 68, p. 1223-1246. Jensen, M# L., 1959* Sulfur isotopes and hydrothermal mineral deposits: Econ. Geol., v. $k9 P. 37^-39^. ________, 1967* Sulfur isotopes and mineral genesis: in Barnes, H. L. (ed.), Geochemistry of hydrothermal ore deposits, Holt, Rinehart, and ¥inston, Inc., New York, p. 143-165. Jones, G. E. and R. L. Starkey, 1957* Fractionation of stable isotopes of sulfur by micro-organisms and their role in deposition of sulfur: Appl. Microbiol., v. 5* P. 111-118. Kajiwara, Y. and J. Date, 1971* Sulfur isotope study of Kuroko-type and Kieslager-type strata-bound massive sulfide deposits in Japan: Geochem. Jour., v. 5* p. 133-150. Kajiwara, Y., 1971* Sulfur isotope study of the Kuroko- ores of the Shakanai No. 1 deposits, Akita, Prefecture, Japan: Geochem. Jour., v. 4, p. 157-181. . 1973a, Chemical composition of ore-forming solution responsible for the Kuroko-type mineraliza tion in Japan: Geochem. Jour., v. 6, p. 141-149. ________, 1973b, A simulation of the Kuroko-type mineraliza tion in Japan: Geochem. Jour., v. 6, p. 193-209. Kaplan, I. R., T. A. Rafter, and J. R. Hulston, i960, Sulfur isotope variations in n ature. 8. Application to some biogeochemical problems: New Zealand Jour. Sci., v. 3, P. 338-3^5. 181 Kaplan, I. R. , K. O. Emery, and S. C. Rittenberg, 1963* The distribution and isotopic abundance of* sulfur in recent marine sediments off* southern California? Geochim. et Cosmochim, Acta, v. 27, p. 297* Kaplan, X. R. and S. C. Rittenberg, 1964, Microbial fractionation of sulfur isotopes: Jour. Gen. Microbiol., v. 34, p. 195# Kaplan, X. R., R. E. Sweeney, and A. Nissenbaum, 1969, Sulfur isotope studies of* Red Sea geothermal brines and sediments: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer—Verlag New York, Inc., p. 474— 498. Kennedy, G. C., 1930, A portion of the system silica- water: Econ. Geol., v. 45, p. 629-653# Ku, T. L., 1969, Uranium series isotopes in sediments from the Red Sea hot-brine area: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer-Verlag New York, Inc., p. 312-524. Ku, T. L., D. L. Thurber, and G. G. Mathieu, 1969, Radio carbon chronology of Red Sea sediments: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer-Verlag New York, Inc., p. 348-367. Lafon, G. M., 1969, Some quantitative aspects of the chemical evolution of the oceans. Unpublished Ph.D. dissertation, Northwestern University. Lambert, I. B. and T. Sato, 1974, The Kuroko and as sociated ore deposits of Japan: a review of their features and metallogenesis: Econ. Geol., v. 69, p. 12 1 5-1 2 3 6. Lawrence, J. R., 1974, Stable oxygen and carbon isotope variations in the pore waters, carbonates, and silicates, sites 225 and 228, Red Sea: Initial Reports, Deep Sea Drilling Project, v. 23, p. 939- 942. Lerman, A., 1967, Model of chemical evolution of a chloride lake— the Dead Sea: Geochim. Cosmochim. Acta, v. 31, P. 2309-2330. 182 Malinin, S. D. and N. X. Khitarov, 1969, Reduction of* sulfate sulfur by hydrothermal conditions: Geochem. Intern., v. 6, p. 1020-1025# Manheim, F. T., 1974, Red Sea geochemistry: Initial Reports, Deep Sea Drilling Project, v. 23, P# 975- 998. Manheim, F. T. , L. S. Waterman, C. C. Woo, and F. L. Sayles, 1974, Interstitial water studies on small core samples, Leg 23 (Red Sea): Initial Reports, Deep Sea Drilling Project, v. 23, p# 955-968. Miller, A. R., 1964, Highest salinity in the world ocean?: Nature, v. 203, p# 590. Miller, A. R., C. D. Densmore, E. T. Degens, J. C. Hathaway, F. T. Manheim, P. F. McFarlin, R. Pockling- ton, and A. Jokela, 1966, Hot brines and recent iron deposits in deeps of the Red Sea: Geochim. Cosmochim. Acta, v. 30, p. 341-359# . 1969, Atlantis II account: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer-Verlag New York, Inc., p. 15-17# Milliman, J. D., D. A. Ross, and T. L. Ku, 1969, Precipita tion and lithification of deep sea carbonates in the Red Sea: Jour. Sed. Petr., v. 39, p# 724-736. Moores, E. M. and F. J. Vine, 1971, The Troodos Massif, Cyprus, and other ophiolites as oceanic crust: evaluation and implications: Phil. Trans. Roy. Soc. London A, v. 268, p. 443-466. Nielsen, H., 1965, Schwefelisotopes im marinen Kreislauf und das S34 der fruneren Meere: Geol. Rundschau, v# 55, P# 160. Neumann, A. C. and C. D. Densmore, 1959, Oceanographic data from the Mediterranean Sea, Red Sea, Gulf of Aden, and Indian Ocean: Woods Hole Oceanographic Institution, Ref. No. 60-2 (unpublished manuscripts). Oana, S. and H. Ishikawa, 1966, Sulfur isotopic fractiona tion between sulfur and sulfuric acid in the hydro- thermal solution of sulfur dioxide: Geochem. Jour., V . 1 , p . 1 * 5 - 5 0 . I83 Ohmoto, H. , 1972, Systematics of* sulfur and. carbon isotopes in hydrothermal ore deposits: Econ. Geol., v. 67, P. 551-578. Ostapoff, F., 1969, A fourth brine hole in the Red Sea?: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer—Verlag New York, Inc., p. 18-21. Phillips, J. D., J. Woodside, and C. 0. Bowin, 1969, Magnetic and gravity anomalies in the central Red Sea in Degens, E. T. and D. A. Ross (eds,), Hot brines and recent heavy metal deposits in the Red Sea, Springer-Verlag New York, Inc., p. 98-113# Puchelt, H., 1969, Sulfur isotope fractionation in minerals: American Report of the Director of the Geophysical Laboratory, Carnegie Institute, Washing ton, D. C., Year Book, v. 68, p. 192-197* Pugh, D. T., 1969, Temperature measurements in the bottom layers of the Red Sea brines: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer-Verlag New York, Inc., p. 158-163# Rafter, T. A. and S. Wilson, 1958, Sulphur isotopic varia tions in nature. Part 5, Sulphur isotopic variations in New Zealand geothermal bore waters: New Zealand Jour. Sci., v. 1, p. 103-126. Robinson, B. W., 1973, Sulfur isotope equilibrium during sulphur hydrolysis at high temperature: Earth Planet Sci. Letters, v. 3, P# 443-450. Robinson, R. A. and R. M. Stokes, 1959, Electrolyte solutions, second edition: London, Butterworths, 559 P* Roedder, E., 1967, Fluid inclusions as samples of ore fluids: in Barnes, H. L. (ed.), Geochemistry of hydrothermal ore deposits, Holt Rinehart, and Winston New York, p. 515-575* Ross, D. A., 1969, Temperature structure of the Red Sea brines: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer-Verlag New York, Inc., p. 148-152. Ross, D. , 1972, Red Sea hot brine area: revisited: Science, v. 175, P. 1455-1957. Ross, D. A. , R. B. Whitmarsh, S. A. Ali, J. E. Boudreaux, R. Coleman, R. L. Fleisher, R. Girdler, F. Manheim, A. Matter, C. Niagrini, P. Stoffers, and P. R. Supko, 1973, Red Sea drillings: Science, v. 179, P. 377- 380. Ryan, ¥. B. F., E, M. Thorndike, M. Ewing, and D. A. Ross, 1969, Suspended matter in the Red Sea brines and its detection by light scattering: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer—Verlag New York, Inc., p. 153-157. Rye, R. 0. and H. Ohmoto, 1974, Sulfur and carbon isotopes and ore genesis: A review: Econ. Geol., v. 69, p. 826-842. Sakai, H., 1957, Fractionation of sulfur isotopes in nature: Geochem. et Cosmochim. Acta, v. 12, p. 150— 169. ________, 1968, Isotopic properties of sulfur compounds in hydrothermal processes: Geochem. Jour., v. 2, P. 29-39. Sakai, H., S. Osaki, and M. Tsukagishi, 1970, Sulfur and oxygen isotopic geochemistry of sulfate in black ore deposits of Japan: Geochem. Jour., v. 4, p. 27-39. Sangster, D. F., 1968, Relative sulphur isotope abundances of ancient seas and strata-bound sulphide deposits: Geol. Soc. Canada, Proc., v. 19, p. 79-91. ________, 1971, Sulphur isotopes, strata-bound sulphide deposits, and ancient seas: Soc. Mining Geol. Japan, spec, issue 3, P. 295-299. Sasaki, A., 1970, Seawater sulfate as a possible deter minant for sulfur isotopic compositions of some strata- bound sulfide ores: Geochem. Jour., v. 4, no. 1, p. 41-51. Sawkins, F. J., 1972, Sulfide ore deposits in relation to plate tectonics: Jour. Geology, v. 80, p. 377-397. 185 Sayles, F. L. and J. L. Bischoff, 1973» Perroraangoan sediments in the equatorial east Pacific: Earth Planet. Sci. Letters, v. 19» p. 330-336. Schneider, E. D. and G. L. Johnson, 1970, Deep-ocean diapir occurrences: Amer. Assoc. Petrol. Geol. Bull., v. 54, p. 2151-2169. Schoell, M. and M. Hartmann, 1973» Detailed temperature structure of the hot brines in the Atlantis XX Deep area, Red Sea: Mar. Geol., v. 14, p. 1—l4. Schoell, M., 1974a, Valdivia VA OI/03 Rotes Meer-Golf von Aden. Hydrographie II. Daten. Stationen VA 01-1 bis VA 01-311s Bundesanstalt f. Boden- forschung, Hannover. ________, 1974b, Valdivia, VA 01/03 Rotes Meer Golf von Aden. Hydrographie III. Daten. Stationen VA 03-322 bis VA 03-636: Bundesanstalt f. Boden- forschung, Hannover. ________, 19759 Heating and convection within the Atlantis II Deep geothermal system, Red Sea. United Nations Geothermal Symposium, San Francisco, California, May 20-29, 1975. Schoell, M., H. Backer, and A. Baumann, 1974, The Red Sea geothermal systems, new aspects on their brines and associated sediments: Intern, assoc, on the genesis of ore deposits (JAGOD), Varna meeting. Schoen, R. and R. O. Rye, 1970, Sulfur isotopic distribu tion in solfataras, Yellowstone National Park: Science, v. 170, p. 1082-1084. Shanks, ¥. C., J. L. Bishoff, and I. R. Kaplan, 1974, Sulfur isotope studies of evaporites and shales from sites 225* 227, and 228 in the Red Sea: Initial Reports, Deep Sea Drilling Project, v. 23, p. 947- 950. Shishkina, O. V. and A. N. Bogoyavlenskiy, 1970, Salt composition of deep waters of the Red Sea: Geokhimiya, no. p. 632-636. Sillitoe, R. H., 1973» Environments of formation of volcanogenic massive sulfides: Econ. Geol., v. 68, p. 1321-1325. 186 Solomon, M. , 1974, Massive sulfides and plate tectonics: Nature, v. 249, p. 821-822* Steiner, A. and T. A. Rafter, 1966, Sulfur isotopes in pyrite, pyrrhotite, alunite, and anhydrite from steam wells in the Taupo Volcanic Zone, New Zealand: Econ. Geol., v. 61, p. 1115-1129* Stoffers, P. and R. Kuhn, 1974, Red Sea evaporites: a petrographic and geochemical study: Initial Reports, Deep Sea Drilling Project, v. 23, P* 821-848. Styrikovich, M. A., 0. I. Nartynova, Z. S. Belova, and V. L. Menshikova, 1968, Electrolytic properties of sodium sulfate in high temperature water: Dokl. Akad. Nauk. SSSR., v. 182, p. 644-646. Swallow, J. C., 1969* History of the exploration of the hot brine area of the Red Sea-Discovery account: in Degens, E. T. and D. A. Ross (eds.). Hot brines and recent heavy metal deposits in the Red Sea, Springer- Verlag New York, Inc., p. 3-9# Tatsumi, T., 1965, Sulfur isotopic fractionation between co-existing sulfide minerals from some Japanese ore deposits: Econ. Geol., v. 60, p. 1645-1659* Thode, H. G. and J. Monster, 1965, Sulfur isotope geo chemistry of petroleum, evaporites, and ancient seas, from: Fluids in subsurface environments: Amer. Assoc. Petr. Geol., Memoir 4. Tooms, J. S. and M. Rugheim, 1969* Additional metal liferous sediments in the Red Sea: Nature, v. 223, p. 1356-1359. Traper, H. G., 1969, Bacterial sulfate reduction in tile Red Sea hot brines: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer-Verlag New York, Inc., p. 263-271* Turner, J. S., 1969, A physical interpretation of the observation of hot brine layers in the Red Sea: in Degens, E. T. and D. A. Ross (eds*), Hot brines and recent heavy metal deposits in the Red Sea, Springer- Verlag New York, Inc., p. 164-173* 187 Upadhyay, H. D# and D. F. Strong, 1973» Geological setting of* the Betts Cove copper deposits, Newfound land: An example of ophiolite sulfide mineraliza tion: Econ* Geol., v. 68, p. 161-167. Vinogradov, A. P., V. A. Grineko, and V. I. Ustinov, 1962, Xsotopic composition of sulfur compounds in the Black Sea: Geochem* Intern*, no* 10, p* 973-997* Voge, H. H* and ¥. F* Libby, 19379 Exchange rates with radio-sulfur: Jour* Amer. Chem. Soc., v. 599 p. 2474. von der Borch, C. C* and R. ¥. Rex, 1970$ Amorphous iron oxide precipitates in sediments cored during Leg 5. Deep Sea Drilling Project: in McManus, D. A. (ed.), Initial reports of the Deep Sea Drilling Project, v* 5, United States Government Printing Office, ¥ashington, D. C* von der Borch, C* C*, ¥. D. Nesteroff, and J. Galehouse, 1970, Iron-rich sediments cored during Leg 8 of the Deep Sea Drilling Project: in Tracey, J. (ed*), Initial reports of the Deep Sea Drilling Project, v. 8, United States Government Printing Office, ¥ashington, D* C. Watson, S, ¥, and J. B. Waterbury, 1969* The sterile hot brines of the Red Sea: in Degens, E* T* and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer-Verlag New York, Inc., p. 272-281* ¥eiss, R. F., 1969, Dissolved argon, nitrogen, and total carbonate in the Red Sea brines: in Degens, E. T. and D. A. Ross (eds.), Hot brines and recent heavy metal deposits in the Red Sea, Springer-Verlag New York, Inc., p. 254-262. ¥hite, D* E*, 1968, Environments of generation of base metal ore deposits: Econ* Geol., v. 63 (4), P. 301- 335. ________, 1974, Diverse origins of hydrothermal ore fluids: Econ* Geol., v. 69, p. 954-973. ¥ood, J. R., 1975* Thermodynamics of brine-salt equilibria I. The systems NaCl-KCl-MgCl2-CaCl2-H20 and NaCl- MgS0.-H20 at 25° C: Geochim. Cosmochim, Acta, v. 39» p. 1147-1163. 188 APPENDIX Computer program 1 1 TURKEY” to calculate distribu tion of dissolved species in the Red Sea brines* 189 101 FORMAT(IX,8F8.4) 105 FORMATIIX,/////) kk3 FORMATI I X ,F10,5 m rH ox p H H O *X R * £ - 1 i r \ m>— » • > • * o mom •*—s o ' »>\ • H • X X 8* * m w o f t ) o h m X ~— ' fa X fe \ m C O H x X •‘H w \ • R w m m X ' ' r w o * • > • * i —1 m • * • * H 0 g £ " » • « o • X X X X ^ m . . m h h s— •>O « a * i m N m H O ►X 5 3 • —• > * • X H X * fn O O * *pt) H ^ » H \ X H X •*X 1 5 r \ m R h m \ m ^ ^<j 'w' • > * »'w' •»*»•» m m^-N m ^ m x oo x m vo o o h • * X • H *H X • • * H HO K m o w \ w mo h * • * h H • H -cf \ C \ i • * H I I VO VO * m r w \ w m R w w w m H • * ' —' f - l • » £ h IW/—' ( x ) ( x ) 'W ^ o m ^ ^ o^'-'0\ £ • < 8- < S H • > g \ g H “N H H H ^ SfoXPn » > f E j X N O I I PS p3 R O w H O m O w H \ f i ^ ^ [s P3 cm m R cm R cm m-a* m m vo - 3- m vo - 3- m vo • % m m mvo • > m h m O o\ H 03 R 9 - 1 • » 53 1 S J * P o 8 -* • % fx l & * O H to 3 X * H W & CM O O H R to * h r t Q P 3 X y —» $ H * O ^ • H • O • » I N - H m i I I P I I H < P to a 5 k p s ^ x m-a- h ON 00 O CM O m m m • • o cm o o I I I I I I o y <; CQ f f l o m m rH mvo on ON^t cm m o 00 00 ON H m cm CM H o H I I CM m-3- P P P H o o o p W 8 W O R R R P o o o o cm m o o o o p p o o o o o o o o P P 53 53 O O N N O O O O o o o R R R R o o H H I S 3 I S 3 R O O O U O O O O O O O cm m P P o o « « Pi R O O CPBCL4=10.**(-1*38) CKCL=10 . **10 . CNACLrrlO • **10 • CHS0=10.**(-1.99) CKS0=10. **(-0.8h) CCAS0=10.**(-2.31) CMGSO=10* ** ( —2 #25) CNAS0=10.**(-1.11) CHC0=:10 . ** (10. 32 ) CH2CO=slO .**(-6.35) CCAC0=10.**(-3.20} CNAC0=10.**(—1.27) CMGC0=10.**(-3.) CMGHC0=10.**(11.22) CCAHC0=10.**(11.58) CHCL=10.**(6.l) DDD=0.0 CLH2S=125.55 CLHS=132.53 CLS=1^6.43 AHsslO . **AH CL¥=-l4. DO k J=1.2 TFE=10.**TPE TCU=10.**TCU TZN—10.**TZN TPB=10.**TPB GFECL3=GO GCUCL2=GO GZNCL2=G0 GPBCL2=G0 GHS10=G0 GHCL=G0 GMGCO=GO GCACO=GO GH2C0=G0 GMGSO=GO GCASO=GO GH2S=G0 GLD=sAL0G10(G0) GLFEC3=GL0 GLCUC2=GL0 GLZNC2=GL0 GLPB C 2=GL0 GLHS 10=sGL0 GLHCL=GLO GLMGCO=GLO GLCACO=GLO GLH2 C O=GLO 191 GLMGSO=GLO GLCASO=GLO GLH2S=GLO GLNACL=GLHCL GLKC L=GLHCL GNACL=GHCL GKCL=GHCL C COMPUTE IONIC STRENGTH XFE3=TFE XCU2=TCU XZN=TZN. XPB=TPB XHS==0# O XS=0.0 XSOrrTSO XHS0=0# O XKS0=0.0 ~XNAS0=0 «O XCOssTCO XHC0=0.0 XNAC0=0# O XMGHC0=0.0 XCAHC0=0# O XNA=TNA XK=TK XCA=TCA XMG=TMG XCL=TCL DO 3 1=1.20 AI=0.5*((XS+XSO+XCO+XCA+XMG)*4•0+(XHS+XHSO+XKSO+ 2 XNASO+ 1XHC0+XNAC0+XMGHC0+XCAHC0+XNA+XK+XCL+XNACL+XHCL)) A12 =XQRT(AI) C DEBYE-HUCKEL CALCULATIONS C CU1 Z=l. A0I=2.5 AG=(-A*Z*AI2)/(1.+A0I*AI2)+BO+AI GLCU1=AG AG=10,**AG GCU1=AG C K 9CL,FECL2,CUCL,ZNCL,AND PBCL Z—10 o A0I=3.0 AG=(-A*Z*AI2)/(1.0+A0I*B*AI2)+BO*AI GLFEC2=AG GLCUCL=AG GLZNCLssAG 192 GLPBCL=AG GLK=AG GLCL=AG AGsslO # **AG GFECL2=AG GCUCLsrAG GZNCLsAG GPBCL=AG GK=AG GCL=AG C MGHCO,CAHCO,HS,HSO,KSO,NASO,HCO,NACO C FECL4,CICL2,CUCL3,ZNCL3,PRCL3,MGOH,FEOH Z=rl.O AOXrr^.O AG= ( -A*Z*AI2 ) / (1. 0+A0I*B*A12 ) +BO*AI GLFEC4=AG GLCIC2=AG GLCUC3=AG GLZNC3=AG GLPBC3=AG GLMGOHs=AG GLFEOH=AG GLMGHC=AG GLCAHC=AG GLHSO=AG GLKSO=AG GLNASO=AG GLHCOrrAG GLNACO=AG GLHS=AG AG=10,**AG GCXCL2s=AG GFECL4=AG GCUCL3=AG GZNCL3=AG GPBCL3=AG GMGOH=AG GFEOH=AG GMGHCO=AG GCAHCO=AG GHS=AG GHSO=AG GKSO=AG GNASOs=AG GHCO=AG GNACO=AG C NA Z r r l * O AOI=4# 5 193 AG= ( -A*Z*AI2 ) / (1. 0+A0I*B*A12 ) -fBO*AI GLNA=AG AG=slG# **AG GNA=AG C H Z=1.0 A0I=9* O AG=(-A*Z*A12) / (1.0+A0I*B*A12(+BO*AI GLH=AG AG=10.**AG GHrrAG C S04 AND ALOH Z=4#0 AOI=4.0 AG-(—A*Z*AX2)/(1♦0+A0I*B*A12)+BO*AI GLSO=AG AG=10.**AG GSO=AG C C03,PB,CICL3,CUCL4,ZNCL4,PBCL4 Z=k.O AOI=4.5 AG= ( —A*Z*A12 ) / (1. O +AOI *B* A12 ) +BO*AI GLPBssAG GLCIC3=AG GLCUC4=AG glznc4=ag GLPBC4=AG GLCO=AG AGaslO. **AG GPB=AG GCICL3=AG GCUCL4=AG GZNCL4=AG GPBCL4=AG GCO=AG C S AND BA AND AL04 z=4.o AOI=5# O AG=(-A*Z*A12)/(1.0+AOI*B*A12)+BO*AI GLS=AG AG=10.**AG GBAnAG GS=AG G CA,PE2,CU2,ZN f FECL Z = 4 . 0 AOI=6 # O AG=(—A*Z*A12)/(1.0+A0I*B*A12)+BO*AI GLFE2=AG 194 GLCU2=AG GLZN=AG GLFECL=AG GLCA=AG AG=XO.**AG GFE2=AG GCU2=AG GZN=AG GFECL=AG GCA=AG C MG Z=4.0 A0Xs=8# O AG= ( —A*Z*A12 ) / (1. 0+A0I*B*A12 ) +BO*AI GLMGssAG AG=10•**AG GMG=AG C FE3 Z=9. AOX=9. AG= ( -A*Z*A12 ) / (1. +A0I*B*A12 ) +BO*AI GLFE3=AG AG=10#**AG GFE3=AG C DEFINE ACTIVITIES AMGOH=GMGOH*XMGOH AFE2=GFE2*XFE2 AFEOH=GFEOH*XFEOH AFECL=GFECL*XFECL AFECL2 =GFECL2 *XFECL2 AFECL3=GFECL3*XFECL3 AFECL4=:GFECL4*XFECL4 ACU1=GCU1*XCU1 ACICL2=GCICL2*XCICL2 ACICL3=:GCICL3*XCICL3 ACUCL=GCUCL*XCUC L ACUCL2 =GCUCL2 *XCUCL2 ACUCL3=GCUCL3*XCUCL3 ACUCL4:=GCTJCL4*XCUCL4 AZNCL=GZNCL*XZNCL AZNCL2=GZNCL2 *XZNCL2 AZNCL3=GZNCL3*XZNCL3 azncl4=gzncl4*xzncl4 APBCL=GPBCL*XPBCL APBCL2 a=GPBCL2 *XPBCL2 APBCL3=GPBCL3*XPBCL3 APBCL4=GPBCL4*XPBCL4 ACU2 =GCU2 *XCU2 195 o co x CO li X co * CO to I ! CO S H . S w X CO Uj co * to O X CO CO * O X CO o X CO o I ! X CO o CO o td UJ X id p X U J 0 J> {> CO CO CO CO o O O O I I I I I I I I U J U J X U J U J U J X X ! > ! > > !> {> UJ * * S O S ! * 0 > > O X X * * ^ co f4 t"* O S O Q O O W O ^ C O CO C0' - ~ to X ^ ^ ^ X co co a a o co + o & o x o X * 0 j> i> * X 0 CO CO CO 0 CO X O O O X + co * * * CO J d O Q Q Q O c o w g o s : + x UJ CO o + UJ X CO o + X o s* CO o + 0 |> |> CO CO CO o o o w u r r r r co to co X X co I I CO I I W M O H I 0 I I CO I I O H X 0 I I !> o co r 0 t r * * O CO F O * I o co 0 * O + O H * » l T ' S> + O X X co X to*^ Ul 0 > CO w f o r o co I X US'— ' > — to uj i * CO o X I X O to UJ to * to X C O 0 I to Q 1 I t 1 0 UJ tr* to UJ co C O I to * X o to S£J O II UJ o H I I X O H H ' * o a to * • x • o x x o t > * * & J O 0 * Q * * !> * I * H X X X X r ^ > o o *so *j X Ci x ^ x x cd |t> O X * ‘ r p o i U S p F J z o l S J s z p K X X X o 0 {> 1 1 {> > o 0 !> {> to I I I I H X I I I I CD I I UJ UJ o o o o {>0 1 1 0 0 X 0 O O O O O O x o 0 k o * x o o i i ( i i i i t ' > I I I I 0 0 0 Q UJ 4* X > i n 'w' \ J m f to • CO H O X * 0 0 g a o u j x o s 0 |> j> t o Z J ^ Q O O O O > US UJ o o o o o o * * * * * o X X X X X * S X O ffi o x 0 > { > t o |> g a o o o & 0 o O O O o X o o o Jt» {t* {t> {s> J^ {s> {^ ts> [t> {t> {> oagpxuJcoS2jxuJNXX {>0 { > C O C O O { > O C O X U I X COCOCOOOII O P H II II u o o o h ii 0 r ii o 0 0 ii II II 0 0 CO II X II N X 0 O 0 0 0 X U J O X X X 52 J W X * X g o co co * * X X I > 0 { > O O X { > X C O X X U ) ----- * C O O * H ( S I X * o x 0 o x w ■' * X * o o O I I I I 0 0 O I I U S ■ ■ o o * O C O C O C O X o o o o X ^ X X X co co o x g o o ‘ > CD l> C O C O C O o o o 0 O 0 x x x *> o x X u CO H o J m f vo G\ 0 CO o + X 2 »> CO o XHSO=RHSO*XSO XNASO=RNASO*XSO XMGSO=RMGSO*XSO XCASO=RCASO*XSO XKSO=RKSO*XSO ALH2 S=RLH2 S +XLS0+GLH2S ALHS=RLHS+XLSO+GLHS ALS=RLS+XLSO+GLS C CARBONATE RH2COs(GCO*(AH**2# O)*CHCO)/(CH2CO*GH2 CO) RHCO=(GCO*AH*CHCO)/GHCO RCACO=(GCO*ACA)/(CCACO*GCACO) RNACO=(ANA*GCO)/(CNACO*GNACO) RMGCO=(AMG*GCO)/(CMGCO*GMGCO) RMGHOO=(AMG+AH*GCO*CMGHCO)/GMGHCO rcahoo=(ac a*ah*gco*ccahco)/gcahco XCO=TCO/ (1. 00+RHC0+RH2C04RCAC0+PNAC0+RMGC0+RMGHC0+ 2RCAHCO) XH200=RH2CO+XCO XHCO=RHCO*XCO XCAHCO=RCAHCO*XCO XMGHC O=RMGHC 0 *XC O XMGCO=RMGCO*XCO XNACO=RNACO*XCO XCACO=RCACO*XCO C CHLORIDE RHCL= ( AH*GCL) / ( CHCL*GHCL) RKCL=(AK*GCL)/(GKCL*CKCL) RNACL=(ANA*GCL)/(GNACL*CNACL) xcl=tcl/(1.O+RKCL+RNACL+RHCL) XKCL=RKCL*XCL XNACL=RNACL*XCL XHCL=RHCL*XCL C MAGNESIUM XMGOH=(XMG*GMG*AOH)/(CMGOH*GMGOH) XMG=TMG-XMGOH-XMGS 0—XMGC O—XMGHC O C HYDROGEN ION EQUILIBRIA IF(M.EQ,2)GO TO 2 XH=AH/GH TH=XH+XHS0+XHC0+2*XH2C0+2*XH2S+XHS+XHCL+XCAHC0+XMGHC0+ 2XMGOH GO TO 1 2 CONTINUE QMGOH=XMGOH/XH XHCL=(AH*ACL)/(CHCL*GHCL) QCAHCO=XCAHCO/XH QMGHCO=XMGHCO/XH QHCL=XHCL/XH QHS=XHS/XH 197 QH2 S=XH2 S/XH QH2 C0=XH2 CO/XH QHCO=XHCO/XH QHSO=XHSO/XH XH=TH/(l«+0HS0+QHC0+2*QH2C0+2*QH2S+QHS+QHCL+QCAHC0+ 2QMGHC0+QMG0H) AH=XH*GH 1 CONTINUE C POTASSIUM XK=TK—XKSO—XKCL C SODIUM XNA=TNA—XNASO—XNACO—XNACL C CALCIUM XCA=TCA—(XCASO+XCACO+XCAHCO) C IRON RLPE2=C LFE2+GLFE2+(O.5*ALW)-ALH—GLFE2 —(O # 2 5*F02) RFE2=10•**RLFE2 RFEOH=(GHE2*A0H*RFE2)/(GFEOH*CFEOH) RFEC L=(GFE3 *ACL)/(GFECL*CFECL) RFECL2=(GFE3*ACL**2)/(GFECL2*CFECL2) RFECL3=(GFE3*ACL**3)/(GFECL3*CFECL3) RFECL4=(GFE3*ACL**4)/(GFECL4*CFECL4) XFE3=TFE/(1.+RFE2+RFEOH+RFECL+RFECL2+RFECL3+RFECL4) XFE2=RFE2 *XFE3 XFEOH=RFEOH*XFE3 XFECL=RFECL*XFE3 XFECL2=RFECL2 *XFE3 XFECL3=RFECL3*XFE3 XFECL4=RFECL4*XFE3 C COPPER RLCU1=C LCU1+GLCU2 + (O.5*AL¥)-ALH-GLCU1-(O.2 5*F02) RCU1=10.**RLCU1 RCICL2=(GCU1*ACL**2*RCU1)/(GCICL2*CCICL2) RCICL3=(GCU1*ACL**3*RCU1)/(GCICL3*CCICL3) RCUCL=(GCU2*ACL)/(CCUCL*GCUCL) RCUCL2=(GCU2*ACL**2)/(GCUCL2*CCUCL2) RCUCL3=(GCU2*ACL**3)/(GCUCL3*CCUCL3) RCUCL4=(GCU2*ACL**4)/(gcucl4*ccucl4) XCU2=TCU/(1.0+RCU1+RCICL2+RCICL3+RCUCL+RCUCL2+ 2RCUCL3+RCUCL4) XCU1=RCU1*XCU2 XCICL2=RCICL2*XCU2 XCICL3=RCICL3*XCU2 XCUCL=RCUCL*XCU2 XCUCL2=RCUCL2*XCU2 XCUCL3=RCUCL3*XCU2 XCUCL4=RCUCL4*XCU2 C ZINC RZNCL= ( GZN*ACL) / ( GZNCL*CZNCL) RZNCL2 = ( GZN*ACL**2 )/(GZNCL2*CZNCL2 ) 198 RZNCL3=(GZN*ACL**3)/(GZNCL3 *CZNCL3) rznciA=(gzn*acl**4)/(GZNClA *CZNClA) XZNsTZN/(1. +RZNCL+RZNCL2+RZNCL3+RZNCIA) XZNCL=RZNCL*XZN XZNCL2=RZNCL2*XZN XZNCL3=RZNCL3*XZN xznciA=rznciA*xzn C PB rpbcl= (gpr*a c l) / ( cpbcl*gpbcl) RPBCL2= ( GPB*ACL**2 . ) / ( CPBCE2*GPBCL2 ) RPBCL3=(GPB*ACL**3*)/(CPBCL3*GPBCL3) RPBCL4=(GPR*ACL**4. ) / (CPBClA*GPBClA) XPB=TPB/(1* +RPBCL+RPBCL2 +RPBCL3+RPBCIA) XPBCL=RPBCL*XPB XPBCL2=RPBCL2*XPB XPBCL3=RPBCL3*XPB XPBClA=RPBClA*XPB 3 CONTINUE PFFE2 =(XFE2/TFE)*100 * PFFE3=(XFE3/TFE)*100. pfcui=(xcui/t c u)*ioo. PFCU2 =(XCU2/TCU)*100. PFZN=(XZN/TZN)*100. pfpb=(xpb/tpb)*1 0 0. PFCL=(XCL/TCL)*1 0 0. pfna=(xna/tna)*1 0 0. PFK=(XK/TK)*lOO. pfmg=(xmg/t m g)*ioo. PFCA=(XCA/TCA)*100. PFSOss ( XSO/TSO ) *100 . pfco=(xco/t c o)*ioo. TFE=AL0G10(TFE) AFE2 =AL0G10(AFE2) AFEOH=AL0G1O(AFEOH) AMG0H=AL0G10 ( AMGOH) AFE3=AL0G10(AFE3) AFECL=AL0G10(AFECL) AFEC L2 a s AL0G10 ( AFECL2 ) AFECL3=AL0G10(AFECL3) AFEC iA ssALOGl O ( AFEC L4 ) TCU=AL0G10(TCU) ACU1=AL0G10(ACU1) ACICL2=AL0G10(ACICL2) ACICL3=AL0G10(ACICL3 ) AGU2 =ALOGIO(ACU2) ACUCL=AL0G10(ACUCL) ACUCL2=AL0G10 fACUCL2) ACUCL3=AL0G10(ACUCL3) ACUClA=AL0G10(ACUClA ) 199 TZN=AL0G10(TZN) A Z N sA LO G lO ( A ZN ) AZNCL=AL0G10(AZNCL) AZNCL2=AL0G10(AZNCL2) AZNCL3=AL0G10(AZNCL3) azncl4=alggio {azncl^ - ) TPB=AL0G10 ( TPB ) APB=AL0G10(APB) APBCL=AL0G10(APBCL) APBCL2 =ALOG10(APBCL2) APBCL3=AL0G10(APBCL3) APBCL4=AL0G10( APBCL**) A W =A L0G 10( AW) TLNA=AL0G10(TNA) tlk=alogio(t k) TLMG==AL0G10 f TMG} TLCA=AL0G10(TCA) TLS0=AL0G10(TSO) TLC0=AL0G10(TCO) TLCL=ALOG10(TCL) AHS10=AL0G10(AHSXO) AS0=AL0G10(AS0) AHS0=AL0G10 fAHSO) AKS0=rAL0G10 ( AKSO ) ACAS0=AL0G10(ACASO) AMGS0=AL0G10(AMGSO) ANAS0=AL0G10(ANASO) ACOssALOGlO ( ACO ) AHC0=rAL0G10 ( AHCO ) AH2C0=AL0G10(AH2CO) ACAC0=AL0G10lACACO) ANAC0=AL0G10(ANACO) AMGC0=ALOGIO(AMGCO) AMGHC 0==AL0G10 ( AMGHCO ) ACAHC 0=AL0G10(ACAHCO) ANA=AL0G10(ANA) AK=AL0G10(AK) ACA=AL0G10(ACA) AMGssALOGlO ( AMG ) AHCL=AL0G10(AHCL) acl=alogio(a c l) AKCL=AL0G10(AKCL) ANACL=AL0G10 ( ANACL) C ION ACTIVITY PRODUCTS OF SALTS CHA=ANA+ACL CCT=ACA+ACO CAN=ACA+ASO CDOL=ACA+AMG+2 *AC 0 C ION ACTIVITY PRODUCTS OF SILICATES DISSOLUTION CQTZ=AHSIO-2*A¥ 200 CKF=AK+AAL+3*AHSXO-4*AW-4*ALH CAB=ANA+AAL+3*AHSIO-4*A¥-4*ALH CANAL=ANA+AAL+2*AHSI0-A¥-4*ALH CANOR=ACA+2*AAL+2 *AHSI0-8*ALH CMUSC=-10*ALH=AK+3*AAL+3*AHSI0 cneph=-4*alh+ana+aal+ahsio CLEUC=-4*ALH-2*A¥+AK+AAL+2*AHSIO CENST=-2*ALH-A¥+AMG+AHSIO CDI0P=-^+ALH-2*A¥+ACA+AMG+2*AHSI0 CTREM=-l4*ALH-8*A¥+2+ACA+5*AMG+8*AHSI0 CTALC=- 6 *ALH- 4 *A¥+3 +AMG+4*AHSI0 CCHYS=-6*ALH+3*AMG+2*AHSIO+A¥ CFORST+- 5 *ALH+2 *AMG+AHSIO CBRUC+-2 *ALH+AMG+2 *A¥ CGXBB=-3*ALH+AAL+3*A¥ C ION ACTIVITY PRODUCTS OF CLAYS CKAOL=—6*ALH+2*AAL+A¥+2*AHSIO CILLT=-8*ALH-2*A¥+0# 5*AK+0.25*AMG+2.3*AAL+3.5*AHSIO C CHLOR=-16 *ALH+ 5 *AMG+2 *AAL+3 *AHSIO+6*A¥ CNASM=-22*ALH-8*A¥+ANA+7+AAL+ll*AHSIO CKSM=-22*ALH-8*AW+AK+&*AAL+11*AHSI0 CCASM=-44*ALH=l6*A¥+ACA+l4*AAL+22*AHSIO CMGSM=-44*ALH-l6*A¥+AMG+l4*AAL+22*AHSIO C ION ACTIVITY PRODUCTS OF SILICATES-----INCONGRUENT DISSOLUTION CKFAB=sAK-ANA CKFMUQ=2 *AK+2 * ( - ALH ) CMUK=2 *AK-2 *ALH- 3 *AW CDITR=2*ACA=AMG—2*ALH CABLEU=ANA+2 *AW-AK-AHSIO CGIBK= 5 *A¥-2 *AHSIO CCHLK=5*AMG+AHSI0+5*A¥-10*ALH CTLANT=2 *AHSIO-5*A¥ C CP=ACU2 *AFE2 +2 *ALS CGAL=APB+ALS CSP=sAZN+ALS CFES=AFE2*ALS CPY=AFE2*((7./4.)*ALS)+(ASO/4.) + ( 2 **ALH)—A¥ ¥RITE(6,445)T,F02,AI,ALH,TH, 2TLNA,ANA,ANACO,DDD,ANASO, 3PFNA,GLNA,GLNACO,DDD,GLNASO, 4TLK,AK,DDD,OOD,AKSO, 5PFK,GLK,DDD,DDD,GLKSO, 6TLMG, AMG, AMGCO, AMGHC O , AMGSO, 7PFMG,GLMG,GLMGCO,GLMGHC,GLMGSO, 8TLCA,ACA,ACACO,ACAHCO,ACASO, 9PFCA,GLCA,GLCACO,GLCAHC,GLCASO, 1TLSO,ASO,DDD,DDD,AHSD,ALH2S,ALHS,ALS, 1PFSO,GLSO,DDD,DDD,GLHSU,GLH2S,GLHS,GLS, 201 2TLC0, ACO, AHCO, AH2C0,DDD, 3PFCO,GLCO,GLHCO,GLH2CO, DDD WRITE(6,443)AW write(6,666)t c l,a c l,a h c l, 1PFCL,GLCL,GLHGL, 2TFE,AFE2,AFEOH, 2PFFE2,GLFE2,GLFEOH, 3TFE,AFE3,AFECL,AFECL2,AFECL3,AFECL4, 3PFFE3,GLFE3,GLFECL,GLFBC 2,GLFEC 3,GLFEC4, 4TCU,ACU1,ACICL2,ACICL3, 4PFCU1,GLCU1,GLCIC2,GLCIC3, 5TCU, ACU2 , ACUCL, ACUCL2, ACUCL3 , ACUCL^, 5PFCU2,GLCU2,GLCUCL,GLCUC2,GLCUC3,GLCUC4, 6TZN,AZN,AZNCL,AZNCL2,AZNCL3,AZNCL4, 6PFZN,GLZN,GLZNCL,GLZNC2,GLZNC3,GLZNCk, 7TPR,APB,APBCL,APBCL2,APBCL3,APBCL4, 7PFPB,GLPB,GLPBCL,GLPBC2,GLPBC3,GLPBC4 WRITE(6,555)CHA,CCT,CAN,COOL, 2CQTZ,CKF,CAB,CANAL,CANOR,CMUSC,CNEPH,CLEUC,CENST,* 2CDI0P,CTREM,CTALC 2,CCHYS,CFGRST,CBRUC,CGIBB, 3CKADL,CILLT,CCHLDR,CNASM,CKSM,CCASM,CMGSM k,CKFAB,CKFMUQ,CMUK,CDITR,CABLEU,CGIBK,CCHLK,CTLANT, 2CCP,CGAL,CSP, 5CFES,CPY WRITE(6,105) IF(M.EQ,2)GO TO k READ(3,101)CLH2S,CLHS,CLS,CHSD,CKSO,CCASO,CMGSO,CNASO READ(5,101)CHCD,CH2C0,CCACO,CNACO,CMGCO,CMGHCO, 2CCAHC0,CHCL READ(5,101)CKCL,CNACL,CMGOH,CLFE2,CFEOH,CFECL, 2CFECL2,CFECL3 READ(5,101)CFECL4,CLCU1,CCICL2,CCICL3,CCUCL,CCUCL2, 2CCUCL3,CCUCL4 READ(5,101)CZNCL,CZNCL2,CZNCL3,CZNCL4,CPBCL,CPBCL2, 2CPBCL3,CPBCL4 READ(5,10l)T,A,B,B0,G0,CLW,PHI,F02 CMG0H=10.**CMGOH CFE2=10,**CLFE2 CFEOHsslO, **CFEOH CFECL=10# **CFECL CFECL2=10.**CFECL2 CFECL3=10„**CFECL3 CFECL4=10* **cfecl4 CCU1=10.**CLCU1 CCICL2=10***CCICL2 CCICL3=10* **CCICL3 CCUCL=10# **CCUCL CCUCL2=10.**CCUCL2 CCUCL3=10.**CCUCL3 202 CCUCL4=10# **CCUCL4 CZNCL=10.**CZNCL CZNCL2=10#**CZNCL2 CZNCL3=10.**CZNCL3 CZNCL4=10.**CZNCL4 CPBCL=10# **CPBCL CPBCL2=10#**CPBCL2 CPBCL3r=10. **CPBCL3 CPBCL4=10.**CPBCL4 CHCL==10.**CHCL CKCL=10.**CKCL CNACL=10.**CNACL CHS0=10.**CHS0 CKS0=10.**CKS0 CCASOsslO . **CCASO CMGS0=10.**CMGSO CNAS0=10.**CNASO CCAHC0=10# **(—CCAHCO) CHC0=10.**(-CHCO) CMGHC0=10.**(—CMGHCO) CH2C0=10#**CH2C0 CCAC0=10.**CCACO CNAC0=10.**CNACO CMGC0=10.**CMGCO M=2 4 CONTINUE 9 CONTINUE END 203

Linked assets

University of Southern California Dissertations and Theses

Conceptually similar

PDF

Geothermobarometric and geochemical studies of two crystalline terrains of the eastern Mojave Desert, USA

PDF

Calcareous nannoplankton from tropical Indo-Pacific deep-sea cores

PDF

The fabrics of deep-sea detrital muds and mudstones: A scanning electron microscope study

PDF

Seismic sequence stratigraphy and structural development of the southern outer portion of the California Continental Borderland

PDF

Natural decay series isotopes in surface waters, bottom waters, and plankton from the East Pacific

PDF

Early Eocene to early Miocene planktonic foraminiferal biostratigraphy of the western Indian ocean

PDF

A gravity and magnetic study of the Tehachapi Mountains, California

PDF

A surface wave study of crustal and upper mantle structures of Eurasia

PDF

Seismic stratigraphic study of the California Continental Borderland basins: Structure, stratigraphy, and sedimentation

PDF

Stable isotopes in live benthic foraminifera from the Southern California borderland

PDF

Deterministic and stochastic modelling of the high frequency seismic wave generation and propagation in the lithosphere

PDF

Tectonics and geochemical exploration for heavy metal deposits in the Southern Gulf of California

PDF

Stratigraphic and sedimentologic analysis of the Monterey Formation: Santa Maria and Pismo basins, California

PDF

Geometry, kinematics, and a mechanical analysis of a strip of the Lewis allochthon from Peril Peak to Bison Mountain, Glacier National Park, Montana

PDF

Geometry and kinematics of the Heart Mountain detachment fault, northwestern Wyoming and Montana

PDF

Spatial and temporal variation of coda Q[-1] in California

PDF

Petrologic and stratigraphic relationships among middle Ordovician limestones from central Kentucky to central Tennessee

PDF

Late Quaternary erosional and depositional history of Sierra del Mayor, Baja California, Mexico

PDF

A sedimentologic history of Ridge Basin, Transverse Ranges, southern California

PDF

Structure of Santa Cruz-Catalina Ridge and adjacent areas, Southern California Continental Borderland from reflection and magnetic profiling: Implications for late Cenozoic tectonics of Southern...

Asset Metadata

Creator Shanks, Wayne Carlton (author)

Core Title Geochemical and sulfur isotope study of Red Sea geothermal systems

Degree Doctor of Philosophy

Degree Program Geological Sciences

Publisher University of Southern California (original), University of Southern California. Libraries (digital)

Tag geochemistry,Marine Geology,OAI-PMH Harvest

Language English

Contributor Digitized by ProQuest (provenance)

Permanent Link (DOI) https://doi.org/10.25549/usctheses-c29-342025

Unique identifier UC11221048

Identifier DP28538.pdf (filename),usctheses-c29-342025 (legacy record id)

Legacy Identifier DP28538.pdf

Dmrecord 342025

Document Type Dissertation

Rights Shanks, Wayne Carlton

Type texts

Source University of Southern California (contributing entity), University of Southern California Dissertations and Theses (collection)

Access Conditions The author retains rights to his/her dissertation, thesis or other graduate work according to U.S. copyright law. Electronic access is being provided by the USC Libraries in agreement with the au...

Repository Name University of Southern California Digital Library

Repository Location USC Digital Library, University of Southern California, University Park Campus, Los Angeles, California 90089, USA