Diss. ETH No. 12029
Deep-Water Renewal in Lake Baikal
A dissertation submitted to the
SWISS FEDERAL INSTITUTE OF TECHNOLOGY
ZORICH
for the degree of
Doctor of Natural Sciences
presented by
Roland Hohmann
Dip!. Natw. ETH
born 22 October 1965
citizen of Ziirich
accepted on the recommendation of
Prof. Dr. D. M. Imboden, examiner
Prof. Dr. Chr. Schar, co-examiner
1997 Contents
Abstract ...... iii Zusammenfassung ...... v TEJHCbl ...... vii Acknowledgements ...... ix
1. Introduction ...... 1 1.1 Lake Baikal ...... I 1.2 How is Deep-Water Renewal in Lake Baikal Accomplished? ...... 2 1.3 Baikal International Center of Ecological Research (BICER) ...... 6 1.4 Outline ...... 6
2. Stability of the Water Column and Neutrally Buoyant Transport in Lakes ...... 8 2.1 Potential Temperature ...... 8 2.2 Equation of State ...... 10 2.3 Local Stability ...... 13 2.4 Potential Density and Quasi-Density ...... 13 2.5 Neutral Surface and Neutral Track ...... 16 2.5.1 Neutral Surface ...... 17 2.5.2 Neutral Track ...... 18 2.5.3 Horizontal Pressure Gradients and Transport along Neutral Tracks ...... 19 2.6 Conclusions ...... 20
3. Processes of Deep·Water Renewal in Lake Baikal ...... 21 3.1 Introduction ...... 22 3.2 Methods ...... 26 3.2.1 Calculation of Salinity from Conductivity ...... 26 3.2.2 Coefficients of Haline Contraction and Equation of State ...... 28 3.2.3 Water Density and Stability of the Water Column ...... 29 3.3 Field Data ...... 31 3.3.1 Density Stratification in Lake Baikal...... 32 3.3.2 Signals in the Bottom Water ...... 34 3.4 Discussion ...... 36 3 .4 .1 Selenga River ...... 3 8 3.4.2 Academician Ridge ...... 46 3.5 Conclusions ...... 51
4. Thermal Bar ...... ,...... 53 4.1 Introduction ...... 5 3 4.2 Spring Thermal Bar ...... 54 4.3 Autumn Thermal Bar ...... 67 4.4 Summary and Conclusion ...... 74 ii
5. Tritium and Noble Gas Analysis: Concepts and Data Processing ...... ? 5 5.1 Geochemical Background ...... ? 5 5.1.1 Tritium ...... 75 5.1.2 Helium ...... 78 5.1.3 Neon ...... 81 5.2 Calculation of the 3H-3He Age ...... 81 5.3 Experimental Aspects and Performance ...... 85 5.3.1 Sampling and Measurement Procedure ...... 86 5.3.2 Calibration ...... 87 5.3.3 Performance ...... 88 5.3.3.1 Errors in Sampling Depth ...... 88 5.3.3.2 Noble Gas Analysis ...... 89 5.3.3.3 Tritium Measurements ...... 92 5.3.4 Overall Performance ...... 94
6. Distribution of Helium and Tritium in Lake Baikal...... 95 6.1 Introduction ...... 96 6.2 Methods ...... 99 6.2.1 CTD Measurements ...... 99 6.2.2 Noble Gas Analysis ...... 99 6.2.3 Calculation of3H-3He Age ...... 100 6.3 Results ...... 101 6.3.1 Vertical Density Stratification ...... 101 6.3.2 Distribution of He, Ne and 3H in the Water Column ...... 103 6.3.3 Helium Isotopes in Hydrothermal Springs ...... 107 6.4 Discussion ...... 109 6.4.1 3H-3He Ages ...... I 09 6 .4. 2 Comparison with CFC-12 Ages ...... I 11 6 .4. 3 Deep-Water Renewal Rates ...... 11 2 6.4.4 Oxygen Consumption Rate ...... 113 6.4.5 4He Flux from the Lake Bottom ...... 114 6.4.6 A Simple Model for Advective Bottom-Water Ventilation ...... 115 6.5 Summary and Conclusion ...... 120
7. Summary and Outlook ...... 122 7.1 Deep-Water Renewal in Lake Baikal ...... 122 7 .2 Ion Budget...... 126 7 .3 Outlook ...... 127
References ...... 12 9
Appendix ...... 135 A 1 CTD Profiles: Sampling Positions and Sampling Dates ...... 13 5 A2 Results of Noble Gas and Tritium Analysis ...... 140
Curriculum Vitae ...... 145 ill
Abstract
This thesis is concerned with the analysis of processes and rates of deep-water renewal in Lake Baikal, Siberia, the world's deepest and largest lake by volume. Despite its great depth, deep-water renewal in Lake Baikal is relatively fast. The relative saturation of dissolved oxygen exceeds 80% in the entire water column. CFC-12 ages calculated from the vertical distribution of chlorofluorocarbon-12 (CFC-12) nowhere exceed 16 years.
The equation of state of Baikal water was determined including the effect of dissolved ions and silicic acid. Based on the theoretical concept of quasi density, the vertical stability of the water column was analysed. Density variations in the lake are extremely small and mostly the result of variations in temperature. However, in the region of the mesothermal temperature maximum, where temperature is close to the temperature of maximum density and the thermal expansivity is very small, variations in salinity are important. Thus, deep- water renewal is likely to occur in regions where water masses with different (8,Sc)- characteristics meet horizontally. From a detailed set of temperature, conductivity and pressure (CTD) measurements, two important sites of deep-water formation were identified: (l) the Selenga Delta between the Southern and the Central Basins; and (2) the Academician Ridge, separating the Central and Northern Basins.
River induced deep-water renewal is important in the Central Basin. In spring, cold and relatively saline water from the Selenga, the major river entering the lake, forms a density plume which reaches the bottom of the basin during April and early May. Due to entrainment of lake water, the plume transports up to about 125 km3 yr-I of water to the deepest part of the basin. Deep-water renewal in the Northern Basin is predominately caused by horizontal mixing of cold and relatively saline water from the Central Basin and warmer and slightly less saline water from the Northern Basin at Academician Ridge. The resulting water mass can sink on either side of the sill. Whereas in the Central Basin the water mass stays at intermediate depth, in the Northern Basin it sinks to the deepest part. In 1995, the sinking water masses resulted in a cold bottom boundary layer with a volume of approximately 80 km3. In the Southern Basin, no deep-water renewing process could be identified so far.
The development of the spring thermal bar establishing at the south-eastern shore of the Central Basin, and of the autumn thermal bar in the Northern basin was analysed. In both regions, vertical exchange is restricted to the uppermost 200 - 300 m. The spring thermal bar at Boldakovo is to a large extent determined by the northerly flow of warm Selenga water at the lake surface along the shore. Since the river water's salinity exceeds the lake's, the temperature gradient across the thermal bar is superposed by a distinct salinity gradient. Therefore, the spring thermal bar is not a pure thermal bar but rather a thermohaline front. iv
The northerly flow of river water along the shore implies that heat is mainly transported parallelly to the shore and just to a small extent towards the central part of the basin. Thus, the thermal bar migrates very slowly offshore. In autumn, a "pure" thermal bar that is entirely determined by the evolution of the temperature stratification establishes in the northern part of the Northern Basin. In contrast to the thermohaline front at Boldakovo, the thermal bar in the Northern Basin migrates fast towards the central part of the basin.
A detailed study of 3ff-3He ages was conducted to investigate deep-water renewal in Lake Baikal's three major basins between 1992 and 1995. Maximum 3ff-3He ages are 14 - 17 yr in the Southern Basin, 16 - 18 yr in the Central Basin and 10 - 11 yr in the Northern Basin. Deep-water renewal rates deduced from volume-weighted mean 3ff-3He ages below 250 m depth are about IO% yrl in the Southern and Central Basins and 15 % yrl in the Northern Basin. In the Southern Basin, the mean 3H-3He age below 250 m increased steadily from 9.6 yr in 1992 to 11 yr in 1995, indicating that deep-water renewal was slight during this time. Advective bottom-water renewal by surface water was estimated from the mass balance of 3He in the lowermost 200 m of each basin. In the Northern Basin, advective bottom-water renewal was between 80 and 150 km3 yrI and in the Central Basin between 10 and 20 km3 yrl. In the Southern Basin, advective bottom-water renewal was practically zero. v
Zusammenfassung
Im Rahmen dieser Doktorarbeit wurde die Tiefenwasseremeuerung im Baikalsee (Sibirien), dem tiefsten und volumenmiissig grossten See der Ertle, untersucht. Trotz seiner enormen Tiefe wird das Tiefenwasser des Sees effizient ausgetauscht. In der ganzen Wassersiiule betriigt die relative Sauerstoffsattigung mehr als 80%. Die aus der vertikalen Verteilung der Fluorchlorkohlenwasserstoffe (FCKW) berechneten Wasseralter sind nirgends grosser als 16 Jahre.
Die Zustandsgleichung des Baikalwassers wurde unter Beriicksichtigung der gelosten lonen und der Kieselsaure bestimmt. Mit Hilfe des theoretischen Konzepts der Quasi-Dichte wurde die vertikale Stabilitiit in der Wassersaule untersucht. Die Dichteunterschiede im See sind sehr klein und meistens durch Temperaturunterschiede bedingt. In der Region des mesothermalen Temperaturmaximums, wo die Temperatur nahe der Temperatur maximaler Dichte ist und der thermische Ausdehnungskoeffizient sehr klein ist, konnen Salinitiitsunter- schiede wichtig sein. Die Tiefenwasserbildung geschieht vorwiegend in Gebieten, in denen Wassermassen mit unterschiedlicher Temperatur und unterschiedlichem Salzgehalt horizontal aufeinandertreffen. Mit Hilfe zahlreicher Temperatur-, Leitfahigkeits- und Druck-Messungen (CTD-Messungen) wurden zwei wichtige Gebiete der Tiefenwasserbildung identifiziert: I. Das Selenga Delta zwischen dem Siid- und dem Mittelbecken und 2. der Academician Ridge zwischen dem Mittel- und dem Nordbecken.
Im Mittelbecken ist der Beitrag der Zufliisse zur Tiefenwasseremeuerung von Bedeutung. Im Friihling bildet das kalte und relativ saline Wasser der Selenga, des grossten Zuflusses des Sees, einen Dichtestrom, der zwischen Mai und anfangs Juni bis an die tiefste Stelle des Mittelbeckens abtaucht. Wegen des Einmischens von Seewasser nimmt das Volumen des Dichtestroms stark zu. Der Dichtestrom transportiert bis zu ca. 125 km3 yrl Wasser an die tiefste Stelle des Beckens. Ausschlaggebend ftir die Tiefenwasseremeuerung im Nordbecken ist die Mischung von kaltem und salinem Wasser aus dem Mittelbecken rnit warmem und etwas weniger salinem Wasser aus dem Nordbecken in der Region des Academician Ridges. Das resultierende Mischwasser kann auf beiden Seiten der Schwelle abtauchen. Im Mittelbecken schichtet es sich in rnittleren Tiefen ein. Im Nordbecken sinkt es bis auf den Grund ab. 1995 verursachte das abtauchende Wasser eine kalte Bodenschicht rnit einem Volumen von ca. 80 km3. Im Siidbecken konnte bisher noch kein Tiefenwasserer- neuerungsprozess nachgewiesen werden.
Die Entwicklung des "Thermal Bars" (thermische Front) wurde im Friihling an der Siidostkiiste des Mittelbeckens und im Herbst im nordlichen Teil des Nordbeckens untersucht. In beiden Regionen ist der vertikale Austausch von Wassermassen auf die vi obersten 200 - 300 m beschrankt. Die Entwicklung des Thermal Bars vor Boldakovo im Friihling wird hauptsachlich durch die warme Selenga bestimmt, die an der W asseroberflache dem Ufer entlang Richtung Norden stromt. Da die Salinitat des Flusswassers hoher ist als jene des Seewassers, wird der horizontale Temperaturgradient von einem Salinitatsgradienten i.iberlagert. Somit handelt es sich nicht um einen echten Thermal Bar, sondern vielmehr eine thermo-haline Front. Die Nordstromung des Flusswassers entlang des Ufers bewirkt, dass der Warmetransport v.a. parallel zum Ufer geschieht und nur in geringem Ausmass vom Ufer weg in Richtung Seemitte. Die thermmo-haline Front bewegt sich deshalb nur langsam vom Ufer weg. Ein eigentlicher Thermal Bar, der ausschliesslich von der Temperaturverteilung bestimmt wird, entsteht im Nordbecken im Herbst. Im Gegensatz zur thermo-halinen Front vor Boldakovo bewegt er sich im Nordbecken schnell vom Ufer weg.
Um die Tiefenwassererneuerung in den drei Becken des Baikalsees zu untersuchen, wurde zwischen 1992 und 1995 eine umfangreiche Studie der 3fl.3He Wasseralter durch- gefiihrt. Die maximalen 3H.3He Wasseralter liegen im Siidbecken zwischen 14 und 17 Jahren, im Mittelbecken zwischen 16 und 18 Jahren und im Nordbecken zwischen 10 und 11 Jahren. Tiefenwassererneuerungsraten, die aus den volumengewichteten 3fl.3He W asseraltem unterhalb 250 m Tiefe berechnet wurden, betragen im Siid- und Mittelbecken ungeflihr 10 % yr I und ungeflihr 15 o/o yr-I im Nordbecken. Im Siidbecken stieg das mittlere 3fl.3He Alter des Tiefenwassers zwischen 1992 und 1995 von 9.6 auf 11 Jahre an. Mil Hilfe der 3He-Massenbilanz in den untersten 200 m jedes Beckens wurde die advektive Bodenwassererneuerung abgeschatzt. Im Nordbecken betrug sie 80 - 150 km3 yr I, im Mittelbecken zwischen 10 und 20 km3 yr!. Im Siidbecken fand praktisch keine advektive Erneuerung der Bodenschicht statt. vii
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Acknowledgements
This thesis was made possible by financial support from the Swiss Federal Institute of Environmental Science and Technology (EAWAG), the Swiss Federal Institute of Technology (ETH) and the Swiss Federal Office for Science and Education (BBW).
During the time of this thesis I enjoyed the support and encouragement of numerous people whom I would like to thank in the following:
Dieter M. Imboden was responsible for awakening my interest in lake physics with his lectures on physical Iimnology. After my diploma thesis, Dieter showed much confidence in me by letting me do my Ph.D. on Lake Baikal. He was at all times a most helpful and encouraging supervisor, continually motivating me to aim for high standards. He also supported me personally in many ways and always impressed me with his open-minded personality.
Christoph Schiir agreed to be co-examiner of this thesis. I thank him for his friendliness and the interest he showed in my work. I certainly enjoyed his uncomplicated and motivating support.
My colleagues Rolf Kipfer, Markus Hofer and Werner Aeschbach-Hertig helped me in the field, in the lab and in the office, and taught me all that is important about helium isotopes and tritium. Besides that, RoK.i put tremendous effort into organising our research activities in Siberia and was a continual source of motivation and a great help. Markus examined the layout of the final manuscript carefully and critically. Werner helped me with the calibration and processing of the data; I appreciated his concentrated and calm attitude in our sometimes crowded and noisy office.
On two trips to Siberia, Michael Schurter helped me to survive by always managing to find nutrients both the solid and the liquid variety. Apart from that, he was responsible for the technical equipment on all the cruises.
Peter Signer gave me a warm welcome in the Noble Gas Laboratory. Rainer Wieler, Heiri Baur, Urs Menet and Stefan Thiirig helped me to make it through several crises in the laboratory with reliable results!
Frank Peeters and Gabriel Piepke came up with an elegant theory to describe the vertical stability and neutrally buoyant transport in cold, deep lakes. After that, finishing my thesis was peanuts! x
The proof-reading of the manuscript was done by Margrit Rupp, Roland LUthi and Vijay Matta. David Livingstone spent days (weeks?) improving and polishing the language of the final text.
With Manuel Gloor, who took part in the 1993 expedition, I shared the discovery of Irkutsk and the friendship with Luda. Andreas Mathieu conducted a first evaluation of the noble gas and tritium measurements during the course of his diploma thesis (Mathieu, 1995). The cover page of this thesis looks so nice because of the experienced work of an anonymous satellite team from the former USSR combined with the skills of Raoul Schaffner. Additional satellite images were provided by David Llewellyn-Jones and Helen L. Le Core. Chemical data were provided by Laura Sigg. Gerrit-Hein Goudsmit helped with the calibration of the conductivity sensor of the CTD probe. Otti Kocsis provided Figure 7.1. Urs Beyerle and Martin Hirt helped measuring the 1995 samples.
A warm thank you goes to my colleagues from the Limnological Institute of the Siberian Division of the Russian Academy of Sciences in Irkutsk: to Nadja Cherepanova who did the Russian translation of the abstract and was always a caring host during our stays in Irkutsk; to Mikhail N. Shimaraev who helped me out with all kinds of information on Lake Baikal; to Tamara Khodzher for providing chemical data; to Mikhail Grachev who works hard to keep the Lirnnological Institute going; to Nick Granin, Andrei Zhdanov, Valentina Domysheva and Ruslan Gnatovsky who helped to organise the expeditions on the lake; and to the crew of the RV Vereshchagin for taking us almost everywhere we wanted to go, and a few places we didn't.
And to close with a few more personal thanks:
My parents, Sylvia and Heinz Hohmann, always encouraged me to study a subject that would invoke a maximum in personal satisfaction. During my thesis work, they helped me with generous financial support.
Barbara Maey showed a lot of interest in my work and gave me much encouragement and moral support. Besides doing some proof-reading, she edited the text and helped with the layout of the final manuscript. I would like to thank her too for just being around. 1. Introduction
1.1 Lake Baikal
Lake Baikal (Fig. 1.1 ), located in the Baikal Rift in East Siberia, is the deepest and largest lake by volume on Earth. Table 1.1 summarises the most important morphometric and hydrographic data on the lake. The lake holds about one fifth of the global inventory of unfrozen fresh surface water (Herdendorf, 1990). Lake Baikal is located in the foothills of the East Sayan Mountains between the Primorskii and Baikalskii mountain ranges in the west, and the Khamar-Daban, Ulan-Burgasy and Barguzinskii mountain ranges in the east. The lake stretches over a distance of 636 km from the south-western to the northern tip and has a maximum width of approximately 80 km. It is divided into three major basins by underwater sills. A sill in the vicinity of the Selenga Delta separates the Southern Basin (max. depth 1432 m) from the Central Basin (max. depth 1636 m). The Academician Ridge, which runs between Olkhon and Ushkanii Islands, separates the Central Basin from the Northern Basin (max. depth 900 m).
Among the 500 inflows to the lake, which together drain a catchment area of approximately 540000 km2, the Selenga (discharge 31.0 km3 yrl), the Upper Angara (8.3 km3 yr!) and the Barguzin (4.4 km3 yrl) are the most important (Shimaraev et al., 1994). The only outflow is the Angara (65.3 km3 yr1), located at the north-western shore of the Southern Basin. 1853 km downstream, it joins the Yenisei River, which eventually flows into the Arctic Ocean (Martin, 1994). The mean residence time of the water in the lake is approximately 350 yr.
Of the lake's 22 islands, Olkhon Island (690 km2) is the largest (Shimaraev, 1994). It delimits the "Maloye More" (Small Sea) at the southern end of the Northern Basin from the Central Basin. The second largest island is Bolshoi Uskanii Island (9.4 km2), one of the 2 archipelago of four islands located off the "Svyatoi Nos" (Holy Cape) peninsula. It is famed as the favorite haunt of the endemic fresh-water seal (Phoca sibirica).
The formation of the Baikal Rift occurred 35 Myr ago. It gave rise to a deep-water basin within the southern depression (30 - IO Myr) and, later on, to a deep-water basin in the North (10- 3.5 Myr: Mats, 1992, in Martin, 1994). The lake acquired the shape it has now during the Late Pliocene and Lower Pleistocene (Logatchev, 1993). The sedimentary record in the Baikal depression is more than 7 km thick and records more than 15 Myr of the lake's history, making it a promising site for paleoclimate studies (Members, 1992). The Baikal Rift is still active today. Examples of this tectonic activity are the numerous hot springs found in the Selenga region and the hydrothermal vents located at Frolikha Bay in the northern part of the Northern Basin (Crane, 1991; Golubev, 1993; Kipfer, 1996).
During the lake's long existence, a unique ecosystem with a high degree of biodiversity has evolved. 1825 animal species and 569 algal species have been reported so far (Martin, 1994). These are minimal values because numerous new species have Table 1.1: Morphometric and hydrographic data been described since. About 54% of on Lake Baikal (Shimaraev et al., 1994). the animal species and 35% of the Altitude 456 ma. s. !. algal species are endemic (Martin, Maximum depth 1636 m 1994). For the gammarid species Mean depth 731 m alone, the rate of endemism is nearly Surface area 31500 km2 98%. Among the numerous endemic Catchment area 540000 km2 species, prominent examples are the Volume 23015 km3 fresh-water seal (Phoca sibirica) and Drainage 61 km3 yr-l the Omul (Coregonus autumnalis mig- Residence time 350 yr ratorius), a very palatable salmonide.
1.2 How is Deep-Water Renewal in Lake Baikal Accomplished?
Despite its great depth, oxygen concentrations in Lake Baikal are remarkably large. Even at the deepest part, the relative saturation of dissolved oxygen exceeds 80%. CFC-12 ages (i.e. the time since a given water parcel was last exposed at the lake surface) determined from the vertical distribution of chlorofluorocarbon-12 (CFC-12), nowhere exceed 16 years (Weiss et al., 1991). Both observations suggest that deep-water ventilation is very efficient and raise the question of how it is accomplished. 3
56
55
Lake Baikal z 54 ::.....
53
52
0 50 100 km
51 104 105 106 107 108 109 110 Longitude [0 E]
Figure 1.1: Map of Lake Baikal showing isobaths at 400, 700 and 1000 m depth. The lake is divided into Southern, Central and Northern basins: SJ, CI and NJ mark the deepest points of these basins.
Water density is a function of temperature, the concentrations of dissolved chemicals (expressed as salinity S) and pressure. Except of the uppermost 100 m, where temperature follows the seasonal development, water temperatures in Lake Baikal are always close to the in situ temperature of maximum density and the thermal expansivity of the water is therefore small. The concentrations of dissolved chemicals are small and rather homogeneous. despite the fact that they differ by more than a factor of two between the major inflows (Votintsev, 4
200
400
600 ! 800 =g. Q 1000 Figure 1.2: Potential temperature in 1200 the deepest part of Lake Baikal on June JO, 1993. The temperature of 1400 maximum density, Tmd· is plotted as thin dotted line. The arrow indicates a surface water parcel sinking to the lake 1600 bottom by free convection (white area) and forced convection (shaded area). 3 3.1 3.2 3.3 3.4 3.5 3.6 3.7 3.8 0 [°C]
1993). Consequently, density variations in the lake are three-dimensional and extremely small, and mostly - but not always - they are the result of the variation in temperature.
The vertical distribution of potential temperature, 8 (in the following, referred to simply as temperature), at the deepest part of the lake (position Cl in Fig. 1.1) on June 10, 1993, is shown in Fig. 1.2. The most important abbreviations used in this thesis are summarised in Table 1.2. The temperature of maximum density, Tmd• is plotted as a thin dotted line in Fig. 1.2. In freshwater at atmospheric pressure, maximum density is reached at 3.98 °C. T md decreases with increasing pressure at a rate of -0.021 K barl. The temperature profile represents the typical winter stratification in Lake Baikal. At the lake surface, temperature is below Tmd· In the uppermost 200 m, the water column is inversely stratified with respect to temperature. At a depth of about 200 m, the water temper.iture reaches its local maximum, the so-called mesothermal temperature maximum (MTM), and crosses the TmJ-line. The temperature of the MTM, TMTM. is approximately 3.6 °C. Below the MTM, temperature is larger than Tmd and slightly decreases with depth. In the lowermost 30 m, temperature decreases markedly and reaches 3.15 °C at the lake bottom.
The low temperature at the lake bottom implies that cold water ( 8 < 3.15 °C) is mixed to the deepest part and compensates for the effect of geothermal heating. Obviously, such cold water is found only above the MTM during winter stratification. Therefore, the question arises 5
how surface water can sink across the Tmd line to the deep part of the lake. The problem is illustrated by the example given in Fig. 1.2. Consider a water parcel at the lake surface the temperature of which has increased by solar irradiation from 3.13 °C to 3.15 °C. Because of its temperature being closer to the local Tmd than the temperature of the water in the uppermost 80 m, the water parcel is denser than the underlying water masses. It will consequently sink by free convection to the thermocline, located a depth of about 80 m depth. In the region of the MfM between depths of 80 m and 310 m (shaded area in Fig. 1.2), the temperature of the ambient water is closer to the local T md than the temperature of the given parcel. The water parcel has to pass through this layer by forced convection in order to reach the deep part of the lake. Below 310 m depth, its temperature is again closer to the local Tmd than the temperature of the ambient water, and therefore it will continue to sink by free convection to the lake bottom.
Table 1.2: Most important abbreviations used in this thesis.
T temperature, referred to as in situ temperature () potential temperature. referred to simply as temperature MTM mesothermal temperature maximum TMTM temperature of mesothermal temperature maximum Tmd temperature of maximum density
In most lakes, partial or complete turnover of the water column - if it occurs is caused by the seasonal cooling and heating of the water, often supported by wind forcing. However, this mechanism is restricted to the upper few hundred meters. In deep water bodies horizontal density gradients are necessary to induce vertical exchange down to the lake bottom. In lacustrine systems such gradients are the result of lateral salinity and temperature gradients which often occur near river inlets or at sills between different basins of a lake (e.g. Aeschbach-Hertig et al. 1996a). In deep freshwater lakes such as Lake Baikal, the non- linearity of the equation of state of water gives rise to two additional phenomena which may induce vertical exchange. (1) The thermobaric instability is linked to the pressure dependence of Tmd· The downward displacement of a water mass across the Tmd'"line causes the thermal expansion coefficient to change sign and results in a local instability. (2) Since water density is not dependent on temperature in a linear way, the mixing of water masses of different temperatures always leads to an increase in the mean density (Cabbeling: McDougall 1987b). The most prominent example is the so-called thermal bar, where two adjacent surface water masses, one warmer, one colder than T md• mix to form water of maximum density which then sinks to greater depth.
A review of the various studies on deep water ventilation in Lake Baikal is given in Sect. 3. l. 6
1.3 Baikal International Center of Ecological Research (BICER)
The Baikal International Center of Ecological Research (BICER) was founded in 1988 by the Limnological Institute of the Siberian Division of the Russian Academy of Sciences, Irkutsk (Russia). Founder members are Russia, Belgium, Japan, the UK, USA. Switzerland joined in 1992. The main objectives of BICER are the coordination of the limnological research activities on Lake Baikal and the protection of 1he lake's unique ecosystem from developments which may lead to irreversible changes.
Switzerland is represented at BICER by the Swiss Federal Institute of Technology (ETH). The Environmental Physics Department of the Swiss Federal Institute of Environmental Science and Technology (EA WAG) has been participating in a project which concentrates on the deep-water renewal in Lake Baikal. This 1hesis forms part of this project. Most recently, the Environmental Physics Department has been taking part in a projectwhich concentrates on the analysis of recent sediments of the lake.
1.4 Outline
In the course of this thesis, the issue of how deep-water ventilation in Lake Baikal is accomplished was approached as follows. (1) Based on chemical data (L. Sigg, unpublished data) and on a detailed set of temperature, conductivity and pressure (CTD) measurements, the equation of state of Baikal water was determined including the effect of ionic salinity and siiicic acid (Hohmann et al. 1997a). (2) A thorough theoretical analysis of vertical stability and neutrally buoyant transport in deep cold lakes was presented by Peeters et al. (1996). (3) Based on these concepts, from a detailed analysis of the one- and two-dimensional distributions of water density processes of deep-water renewal were identified by Hohmann et al. (1997a). (4) Deep-water renewal rates were determined from 3ff.3He ages calculated from a detailed set of helium, neon and tritium concentrations (Hohmann et al. l 997b).
The data included in this thesis were obtained during the course of four cruises on Lake Baikal organised by BICER:
5 - 15 July 1992; • 18 May- 26 June 1993; 21 October - 14 November 1994; 11 May -4 June 1995. 7
During these expeditions, more than 600 CTD casts were taken, and 281 water samples for the analysis of helium and neon isotopes, and tritium were collected. An overview of the data collected is given in Appendices Al and A2.
In detail this thesis is set up as follows. Section 2 gives a summary of the theoretical concepts to describe stability and neutrally buoyant transport in freshwater lakes (Peeters et al. 1996). The equation of state is formulated in terms of salinity, pressure and either in situ temperature or potential temperature. The concepts of quasi-density and neutral tracks are introduced and compared with the concepts of potential density and neutral surfaces, respectively (Sects. 2.4 and 2.5).
Section 3 consists of a paper accepted to be published in Limnology and Oceanography, in which processes of deep-water renewal are described (Hohmann et al., I997a). Based on chemical data and on a detailed set of CTD casts, the equation of state of Baikal water is determined including the effect of ionic salinity and silicic acid (Sect. 3.2). Deep-water formation at the Selenga Delta between the Southern and the Central Basins, and the Academician Ridge between the Central and the Northern Basins are discussed in detail in Sect. 3.4.
In Section 4, observations of the spring thermal bar occurring at the south-eastern shore of the Central Basin (Sect. 4.2) and of the autumn thermal bar in the Northern Basin (Sect. 4.3) are presented.
Section 5 offers an introduction to the noble gas tritium and analysis. The geochemical background of helium, neon and tritium is summarised in Sect. 5.1, and the calculation of 3H- 3He water age is described in Sect. 5.2. An overview of the experimental aspects and of the performance of the measurements achieved is given in Sect. 5.3.
Section 6 discusses the distribution of helium and tritium in Lake Baikal. It consists of a paper accepted for publication by Journal of Geophysical Research (Hohmann et al., l 997b ). In Sect. 6.3 the results of helium, neon and tritium measurements of samples from the lake are presented and the 3Hef4He ratio of the terrigenic helium component injected into the water through the sediment-water interface is estimated from measurements of water from hydrothermal springs in the vicinity of the Northern Basin. In Sect. 6.4, 3H.3He ages are discussed, and deep-water renewal as well as oxygen consumption rates, and the 4He-flux from the sediment are calculated. Furthermore, a simple model is presented to quantify advective bottom-water renewal rates in the lake's three basins.
The thesis ends - in Section 7 - with a brief resume drawn from the most important results and with an outlook on possible future research. 2. Stability of the Water Column and Neutrally Buoyant Transport in Lakes
Theoretical concepts to describe the stability of the water column and neutrally buoyant transport in deep, cold freshwater lakes were developed by Peeters et al. (1996). These concepts have been successfully used to interpret the CTD data from Lake Baikal (Sects. 3 and 4). In the following, the most important theoretical aspects will be outlined. The reader is referred to Peeters et al. ( 1996) for details.
2.1 Potential Temperature
Consider a water parcel which is moved isentropically, i.e. without exchanging heat or mass, from depth z to a reference depth Zr. Due to the compressibility of water, the density and temperature of the parcel Will change along its path. The potential temperature e(z, Zr) of the parcel at depth z with respect to Zr is defined as the in situ temperature T of the parcel at the reference depth Zr (e.g., Gill 1982): z, e(z,z,) = T(z)- I I'z [e(z,z'),S(z),p(z' )] dz', (2.1) z with
e(z,z') = e[T(z),S(z),p(z),p(z' )], where Tis the temperature, S the salinity (mass of dissolved solids per mass of solution), and p the pressure at depth z. The z-coordinate is chosen positive upwards. I'z is the adiabatic lapse rate in a vertical water column, i.e. the rate of temperature change due to adiabatic compression (see Gill 1982):
dT) = ga T, I'z(T,S,p) = ( (2.2) dz isen Cp 9 with
g: acceleration due to gravity; a(T,S,p) = _]._(Bp) thermal expansion coefficient at constant S and p; p ()T S,p p(T,S,p): density of water; cp(T,S,p): specific heat of water at constant pressure; T: in situ absolute temperature.
The subscript isen stands for isentropic transport. From the definition of potential temperature and salinity it follows that 9(z.z,) and S are constant under isentropic transport. In limnology the reference depth is commonly chosen to be at the Jake surface ( z, = 0). In the following, the reference depth is kept constant and we use the simple notation 9(z) = 9(z,z,).
Generalised to three dimensions the definition of potential temperature is:
x, 8(x,xr)= T(x)- f J18(x,x'),S(x),p(x')]dx' (2.3) x where xis the coordinate vector. 8(x,xr) is the potential temperature of a water parcel with respect to the reference position x ,, and the general adiabatic lapse rate r = (aT/()p);,,n ·(dp/dx). Since only pressure depends on the integration variable x', the integral is independent of the path along which the water parcel is transported (note that 9(x,x') = e[T(x),S(x),p(x),p(x')J. Therefore, (;l(x,x,) depends only on the characteristics of the parcel at x. i.e., on T(x), S(x),p(x), and on the reference pressure p(xr) Thus, for a fixed reference depth, a change in potential temperature can be written as:
d8=:;-ael dT+:;-ael dS+ ael.'l... dp. (2.4) aT S,p dS T,p vp T,S
During isentropic transport, 9 and S remain constant ( d8;,,,, = dS;sen = 0 ). Since (dp)i.m• = dp, Eq. 2.4 can be written as
OTaej (dT)isen =- ael dp. (2.5) S,p a'P T,S
Combining Bqs. 2.4 and 2.5 yields
d8= ~el (dT-(dT)1,.,,]+ ~el dS. (2.6) oT S,p dS T,p 10
In situ temperature and potential temperature at the deepest part of Lake Baikal on June 10, 1993 are shown in Fig. 2.la. The temperature of maximum density (Tmd) is plotted as a dotted line. In freshwater at atmospheric pressure, maximum density is reached at 3.98 °C. Tmd decreases with increasing depth at a rate of 0.021 K barl. The difference between 6 and T depends on the adiabatic lapse rate I' and therefore on the absolute value of the thermal expansion coefficient a (Eq. 2.2). In the entire water column of Lake Baikal, the water temperature is close to Tmd and the absolute values of a and I' are close to zero. The difference of 6 and T increases with depth, but nowhere does it exceed 0.015 "C. Larger differences between 6 and T are encountered in warm, deep lakes, e.g. in the East African Lakes (see Wuest et al., 1996), where the water temperature is much higher than Tmd and the values of a and I' are large.
0
...... , ...... ~ 500 ] 1- 1000 0
1500
3.1 3.2 3.3 3.4 3.5 3.6 ·1 0 1 2 3 4 5 1 Temperature[°CJ u[10'" K ]
Figure 2.1: (a) Profiles of in situ temperature T (solid line) and potential temperature O(dashed line), and (b) thennal expansion coefficient a. The profiles were measured at the deepest part of Lake Baikal on June 10, 1993. The dotted line in (a) represents the temperature of maximum density, Tmd·
2.2 Equation of State
The equation of state can be formulated in terms of in situ values of salinity S, pressure p and either water temperature T or potential temperature 6.
p(6,S,p) = p(T,S,p). (2.7) 11
In the following, functions of 8, Sand p are marked with '""". At any fixed location, p and p have the same numerical value.
According to Eq. (2. 7) density variations can be written in terms of either Tor e:
dp= aAI~p dO+;a·1 dS+;a'/ dp=dp=...J!..a I dT+;a I dS+;a j dp.(2.8) 08 S,p dS (J,p Variations in density due to temperature, salinity and pressure changes are described by the thermal expansion coefficient, by the haline contraction coefficient and by the compressibility (see Gill 1982). Depending on the equation of state, these coefficients are I a(T,S,p) = --()T()pl , (2.9a) p S,p I ()pl fJ y(T,S,p) = -j}I dpl , (2.9c) P Pr.s or a(A e,S,p) = --;:I aplae , (2.lOa) p S.p A I ()pl {J(e,s,p)=-;:as . (2.lOb) p 9,p y(• e,S,p) =-;:j} I ap' . (2.lOc) p 'P o.s Both sets of coefficients are related by: 1 a=a-A (ae)- , (2.lla) ()T S,p • (ae) (2.llb) fJ =fJ + a as r.p' r=r+a-. A(ae) · (2.llc) dp T,S 12 Substituting Eqs. (2.9) and (2. l 0) in Eq. (2.8), the equation of state becomes either (2.12a) or p-'dp =-adT+j3dS+ ydp. (2.l2b) Since (}and S remain constant during isentropic transport, from (2.12) it follows that: (2.13) Fig. 2. lb illustrates the thermal expansion coefficient a calculated using the polynomials of Chen and Millero (1986) for the profile shown in Fig. 2.la. Due to its pressure dependence (Eq. 2.10), the relative change of a with depth is large, although temperature variation in the entire water column is less than 0.5 K. In the 200 m surface layer where T < Tmd, a< 0, i. e. a temperature increase causes a density increase. At about 200 m depth, water temperature is equal to Tmd and a passes through zero and changes sign. Below 200 m, T > Tmd and a> 0, i.e. density decreases with increasing temperature. The variation of a with pressure is given in Table 2.1 for temperatures close to Tmd and S = 0.1 g kg-I (a typical salinity for Lake Baikal). Note that the difference between the numerical values of a and a is negligible. Table 2.1: Thermal expansion coefficient a as a function of 8, S and p calculated using the polynomials of Chen and Millero (1986). Salinity is set to 0.1 g kg-1 which is a typical value for Lake Baikal. p a(9=3°C) a((:) =3.5°C) a(9=4°C) 0:(9=4.5°C) a(9=5°C) Tmd fbarl [10·!0·6 K·11 1°Cl 0 -15.6 -7.4 0.6 8.5 16.3 3.96 10 -12.4 -4.3 3.7 11.6 19.3 3.76 20 -9.1 -1.1 6.8 14.6 22.3 3.56 30 -5.9 2.1 9.9 17.6 25.2 3.36 40 -2.7 5.2 13.0 20.7 28.2 3.16 50 0.6 8.4 16.1 23.7 31.2 2.95 13 2.3 Local Stability The stability of a water column is determined by its vertical density profile. A water column is locally stable if a fluid parcel that is displaced isentropically from its initial position by an infinitesimal vertical distance dz experiences a restoring force. This is equivalent to the condition that the density change of a parcel due to an infinitesimal isentropic displacement must exceed the density change in the ambient water over the same distance dz: dp) - dp >0. (2.14) (dz isen dz Note that in Eq. (2.14) p can be replaced by p since {>( x, y, z) = p( x, y, z). Multiplying Eq. (2.14) with g ·p- 1 and using Eqs. (2.12) and (2.13) yields the square of the Brunt-Vliisala frequency (Gill 1982): NzA 2( O,S,p ) =-;;-g [(di>)- . -dp] g(Ad() a- {J-.·dS) (2.15) P dz isen dz dz dz From Eqs. (2.14) and (2.15) we conclude that a water column is locally stable if&;> 0. For Ni 0 it is neutrally stable and for Ni < 0 it is locally unstable. The Brunt-Vliisiilli frequency can also be expressed in terms of T and S ( Millard et al. 1990): Ni(O,S,p)=Ni(T,S,p)=g_((dp) - dp)=g(a[dT (2.16) P dz isen dz dz +r]-Pds)dz where Eqs. (2.6) and (2.11) have been employed. 2.4 Potential Density and Quasi-Density Potential density Ppo1(z,z,) is commonly used by oceanographers to study the vertical stability of the water column. It is defined as P,,or = p[ 0( z. z,), S( z), p( z,)]. (2.17) Note that in Eq. (2.17) the density function p(T, S, p) (and not f>( (), S, p)) is evaluated for 0( z. z,). It can be rewritten as an integral equation similar to Eq. (2.1 ): 14 z, Ppo1(z,z,) = p(z)- J P(z,z')dz', (2.18) z where the adiabatic lapse rate for density, '¥, is given by: (2.19) Eq. (2.18) refers to a water parcel that is moved isentropically from its initial depth z to a reference depth z,. The density of the water parcel at the reference depth is its potential density. Eq. (2.19) is a generalisation of the commonly used adiabatic density lapse rate -(dp!dz\,.. = '¥(z,z) that has been introduced in Eq. (2.14). Inserting Eq. (2.19) into (2.18) yields: zJ'(dp[e(z.z'),S(z),p(z')]) d, Ppot P ()z + d, z z Z is en (2.20) =p(z)+ f dp[ 9(z,z'~~(z),p(z' )] dz' = p( 9(z, z,),S(z),p(z,)} z z which demonstrates the agreement between Eqs. (2.17) and (2.18). In Eq. (2.20) the subscript isen has been dropped since only pressure p depends on the integration variable z' (note that 9(z,z')= e[T(z),S(z),p(z),p(z')]). Multiplication of the vertical gradient of Ppot with -gp~!1 yields: - - g -d-dppot = g( a ( 9(z.z,),S(z),p(z,) lde-dz -/3 ( 9(z,z,),S(z),p(z,) ldS)- . (2.21) ~ z dz Eq. (2.21) looks similar to the definition of the square of the Brunt-Vaisalli frequency Eq. (2.15). However, in Eq. (2.21) thefunctions a and /3 are used instead of a and p. The arguments of these functions are different in Eqs. (2.21) and Eq. (2.15): in Eq. (2.15), a and p are evaluated at the local depth z; in Eq. (2.21 ), a and /3 are calculated at the reference pressure p( z,). Close to the reference depth, the use of the different functions for the thermal expansion and haline contraction coefficients (a and P; a and /3) causes only a very small disagreement between Eq. (2.21) and (2.15). For z z,, the respective coefficients are identical (see Eqs. 2.1 and 2.11) implying that the arguments in Eqs. (2.21) and (2.15) are the same and ii}: -(g/ppo1 )(dppo1 /dz). However, because of the pressure dependence of the thermal expansion coefficient (Tab. 2.1), Eq. (2.21) can be very different from f:I; far from the reference depth. The potential density calculated from the CTD 15 profile shown in Fig. 2.1 is illustrated in Fig. 2.2. Below 300 m depth, the potential density decreases with increasing depth, suggesting that the water column is vertically unstable. However, from the temperature profile (Fig. 2.la) it becomes clear that the stratification is stable, especially since the distribution of salinity hardly influences the vertical stability (see Sect. 3.3.1). In the deep water, temperature is below 4 °C but above the local value of T md• and the temperature decrease with increasing depth has a stabilising effect (a> 0). Remember that by calculating the potential density the imaginary water parcel is moved from the deep waters across the Tmd curve to the reference depth Zn chosen to be at the lake surface, and a and the gradient of Ppot change sign. Thus, in a deep cold lake such as Lake Baikal the vertical gradient of potential density may lead to a wrong interpretation of the vertical stability of the water column. If the water temperature is far from the temperature of maximum density (as is the case in the East African Lakes, e.g. in Lake Malawi, see Wiiest et al. 1996), or if the water does not have a density anomaly (e.g seawater), the relative change of a with pressure is small and the gradient of potential density can provide a good approximation to the buoyancy frequency. In oceanography, several different reference depths are employed to improve the range over which the concept of potential density remains a good approximation. 0 ...... Figure 2.2: In situ density (solid ·- line), potential density (dashed line) \ I and quasi-density (dotted line) 500 ~ calculated from the CTD profile 1\ measured at the deepest part of Lake ] I \ Baikal on June JO, 1993. For the I \ calculation of the different densities, 1000 the effect of silicic acid has been g I \ included. Note the very different axis I \ scales in the two graphs. 1500 I \ •.. 0.045 0.05 1 3 5 7 3 Density - 1000 [kg m" ] 16 Quasi-density pqua ( z. z,) was introduced by Peeters et al. ( 1996) as an alternative to the concept of potential density. This new quantity allows the vertical stability to be assessed in any system, whether fresh or salty, cold or warm. Quasi-density is defined as z, Pqua(z,z,) = p(z)- f 'P(z',z')dz', (2.22) where the function 'Pis given by Eq. (2.19). The difference between the two concepts potential density and quasi-density lies in the argument of the adiabatic density lapse rate. In Eq. (2.18), 'P(z,z') refers to the adiabatic density lapse rate of the water parcel that has been moved isentropically from depth z to depth z', while in (2.22), 'P(z' ,z') is the local adiabatic density lapse rate of the background field at depth z'. Thus, potential density is a property of a specific water parcel, while quasi-density depends on the distribution of (J, S, and p in the water column. At the reference depth z,, quasi-density and in situ density are equal. By taking the negative vertical gradient of Pqua one obtains the stability condition Eq. (2. 14): - dPqua =- dp+.!!:_[J'P(z',z')dz']= - dp - 'P(z.z) (dp) - dp (223) dz dz dz z dz Therefore, quasi-density can serve as a means of determining the local stability of the water column. The water column is locally stable if Pqua increases with depth (i.e. if Pqua decreases with z). Multiplication of the negative gradient of Pqua with gp-1 yields the square of the Brunt Vliislila Frequency N; (see Eq. 2.16). Note that, as for N}, quasi- density provides only local information on vertical stability. Quasi-density calculated for the profile measured on June 10, 1993 is shown in Fig. 2.2 as a dotted line. In contrast to potential density, quasi-density increases steadily with increasing depth, indicating that the water column is stably stratified. 2.5 Neutral Surface and Neutral Track What are the directions along which water parcels can be displaced buoyancy-free in a two- or three-dimensional (0, S, p)-field? The concepts of potential density and quasi- density provide only a tool for the analysis of the local vertical stability of a water column. However, in many cases, deep-water renewing processes, e.g. the plunging of a river plume or a density current across a sill, are two- or three-dimensional. The analysis of isentropic transport of water parcels in a two- or three-dimensional ( 0, S, p )-field requires additional tools: the neutral surface and the neutral track. 17 2.5.1 Neutral Surface In a two- or three-dimensional ( e, S, p )-field the direction in which a water parcel can be isentropically displaced from its equilibrium position over an infinitesimal distance without experiencing a restoring force is defined as follows (McDougall, 1987a): (2.24) The isentropic density variation of the parcel being displaced in this direction over an infinitesimal distance is equal to the in situ density difference of the ambient water. In a two-dimensional (e, S, p)-field, Eq. (2.24) defines a curve. In a three-dimensional (e, S, p)- field, the ensemble of directions of buoyancy-free displacements defines a neutral surface (McDougall 1984,I987a). In the following, the discussion will be restricted to the (x,z) plane. The projection of the neutral surface onto the (x,z) plane defines a local coordinate system ( 11,(;), where T/ is the tangent and (; is perpendicular to the neutral surface (Fig. 2.3a). From Eq. (2.24) follows: (2.25) The angle 8 between the neutral surface and the (x.z) plane is given by: a-, de - /3-'dS '2 8 = arctan - , j{J , ::S l= arctan(- ~~ ). (2.26) [ a- - /3- ' dz dz where, corresponding to Eq. (2.15), fl; is defined as: .N; = ~[( dp) _ dp] = 8(a de /3-.-dS) (2.27) p dx isen dx dx dx Locally, the direction of the vector N2 =(N';,N;) is perpendicular to the direction of the neutral surface, i.e. it points in the (;-direction of the local coordinate system. By defining .N; accordingly, the above discussion can easily be extended to three dimensions. Although N'; and N; are both components of the vector field N2 and mathematically equivalent to N;, only the z-component has the physical meaning of the square of a stability frequency, since gravity as the restoring force acts only in z-direction. 18 Figure 2.3: Neutral surface and neutral track of a water parcel. The neutral surface (thick solid line) defines the direction along which an infinitesimal isentropic displacement of a water parcel is buoyancr-{'ee. lt is perpendicular to the stability z vector N The projection of the neutral surface onto the (x,z) plane defines a local coordinate system ( q,(;) where 11 is the tangent to the neutral 2 surface and ' is parallel to N . The neutral track (dashed line) defines the path along which a water parcel can move buoyancy-free from its initial ' ' 11 • position over a finite distance. Along the neutral track the density of the water parcel is equal to the ,______.. x density of the ambient water. 2.5.2 Neutral Track As was pointed out by McDougall (1987b), only infinitesimal displacements along the neutral surface are buoyancy-free. The path along which a water parcel can move isentropically from its equilibrium position over a finite distance without experiencing any buoyancy is called its neutral track (Peeters et al., 1996). The neutral track is a combined property of the density field and of the characteristic properties of the selected parcel, i.e. ()pa and Spa. Along the neutral track the density change of the water parcel must equal the density change of the ambient water: (2.28) where Ppa =p[6pa,Spa•P(x,y,z)]. Everywhere along the neutral track the difference between the density of the local ( fJ, S, p )-field and the density of the water parcel is zero. If Eq. (2.28) is multiplied by fr 1 and X is the tangent to the neutral track, it follows from Eqs. (2. l l) and (2.13) that: A dp A dfJ R dS A dp r pa dz = - a dx +/.I dx + r dx (2.29) A dfJ ll dS A A dp =>a- - 1.1-+rr -rJ- 0, dz dz pa dx with ypa = r[ ()pa, Spa• p(x, y,zJ]. Analogous to ii; and ii; we define: 19 Mx2 = ~[(ddpxpa"' ). _ dp]dx =g[, adx d8 - p' dSdx +(Ypa, y')ddxp] P zsen (2.30) A d8 p' dS A g a- - -+(r a y)dp] [ dz dz P dz The vector M2 (M;,M'ff) is perpendicular to the neutral track in the (x,z) plane. By defining M; corresponding to M;, the above definitions can be extended to three dimensions. Each water parcel has its own neutral track, which can differ from that of a neighbouring water parcel. Neutral tracks of different parcels can cross. The topology of a neutral track is detennined by 8pa and Spa of the parcel and the particular (8, S, p)-field. A neutral track can be a volume, a surface, a line or can reduce to a point. Although, according to Eq. (2.28), the density of the water parcel is equal to the density of the background field everywhere along the neutral track, a water parcel that moves along its neutral track may generate an instability. This is because two water masses having different temperatures and salinities but equal densities may have different compressibilities entering the stability condition Eq. (2.14) through Eq. (2.13). 2.5.3 Horizontal Pressure Gradients and Transport along Neutral Tracks The theoretical concepts described above lead to nontrivial solutions only if horizontal density and pressure gradients exist at least at some locations in the water body. Otherwise fli, N;, M;, and M; would all be zero (see Eqs. (2.27) and (2.30)), implying that neutral surfaces, neutral tracks and isopycnals (the isolines of potential density) are horizontal. However, if horizontal pressure gradients exist, transport along a neutral track is not force-free, even though it is buoyancy-free. Consequently, work is required to transport a water parcel along its neutral track in the direction of the horizontal pressure gradient. The work per unit mass, W, is given by the following integral along the neutral track: 1 W= JF•dr= J-- dp(x')dx, (2.31) m 111Ppa( x') dx' 20 where F is the force acting on the parcel along the infinitesimal distance dr, and n t denotes the neutral track. If transported in the direction of the negative horizontal pressure gradient, the water parcel will gain kinetic energy along its neutral track. This must be the preferred direction along which the water parcel moves. However, the parcel may leave its neutral track if it reaches a position where it causes a vertical instability, although its density corresponds to the density of the surrounding water. Such situations will be discussed in Sects. 3 and 4. 2.6 Conclusions The stability analysis of the water column in deep, cold freshwater lakes requires concepts that deal with subtleties related to the equation of state of water close to the temperature of maximum density. Potential density, a common hydrographic tool employed by oceanographers, may yield erroneous results when used to analyse local stability in freshwater systems. To overcome this deficiency the so-called quasi-density was introduced by Peeters et al. (1996). The vertical derivative of quasi-density is proportional to the square of the Brunt-Vaisiilii frequency N'J: and is particularly useful in determining an optimal spatial resolution for calculating N'J:. Quasi-density has turned out to be an excellent parameter for the one-dimensional analysis of the vertical stability of the water column in Lake Baikal. Neutral surfaces and neutral tracks are two complementary concepts useful to describe neutrally buoyant transport in a two- or three-dimensional (8, S, p)-field. The former was introduced by McDougall (1984, 1987a) and defines the direction along which a water parcel can be displaced buoyancy-free from its initial position over an infinitesimal distance. The latter was introduced by Peeters et al. (1996) and describes the path along which a water parcel can move buoyancy-free from its equilibrium position over a finite distance. The directions of neutral tracks and neutral surfaces differ only if the structure of the density field depends on the spatial variation of both temperature and dissolved solids. If a water parcel is isentropically displaced from its equilibrium position over only an infinitesimal distance, the compressibility of the water parcel is equal to that of the surrounding water R. HohmannO, R. Kipfer!}, F. Peetersl), G. Piepkell, M. N. Shimaraev2), and D. M. Imboden I) !)swiss Federal Institute of Technology (ETH) and Swiss Federal Institute of Environmental Science and Technology (EA WAG), CH-8600 Diibendorf, Switzerland 2lLimnological Institute of the Siberian Division of the Russian Academy of Sciences, Irkutsk 664033, Russia (accepted by Limnology and Oceanography) Abstract Deep-water renewal in Lake Baikal (Siberia), the world's deepest lake and largest lake by volume, is relatively fast. Water age calculated from tritium and helium as well as from CFCs does not exceed 19 years. Relative saturation of dissolved oxygen typically exceeds 80%. The equation of state of Baikal water was determined including the effect of dissolved ions and silicic acid. Based on nearly six hundred CTD profiles taken between 1993 and 1995, two mechanisms of deep-water mixing were identified. (1) In spring, cold and relatively saline water from the Selenga, the major inflow to the lake, forms a density plume which reaches the bottom of the Central Basin during April and early May. Due to entrainment of lake water the plume transports about 125 km3 of water per year to the deepest part of the basin. Later in spring, the river water forms the thermal bar observed along the eastern shore. There are indications that parts of the Selenga are also plunging to the deep part of the Southern Basin. (2) At Academician Ridge, separating the cold and "saline" water of the Central Basin from the warmer and slightly less saline water of the Northern Basin, horizontal mixing results in a water mass which can sink on either side of the sill. Whereas in the Central Basin the water mass stays at intermediate depth, in the Northern Basin it sinks to the deepest part. More detailed data are needed to quantify this density flux. No indication of a wind-induced thermobaric instability was found. 22 3.1 Introduction Lake Baikal (Siberia) is the deepest (1632 m) and largest lake by volume (23015 km3) on Earth (Shimaraev et al. 1994). It holds roughly 20% of the global fresh liquid surface water. The lake is located in the great Baikal Rift zone of eastern Siberia. It is divided by underwater sills into three main basins (Fig. 3.1): the Southern Basin (max. depth 1432 m), the Central Basin (1632 m), and the Northern Basin (897 m). Main inflows are the Selenga, the Upper Angara and the Barguzin. The only outlet is the Angara. Despite its great depth, relative saturation of dissolved oxygen exceeds 80% in the entire water column. Based on the vertical distribution of chlorofluorocarbons (CFCs), Weiss et al. (1991) have shown that the water age is nowhere greater than 16 years. Water age is defined here as the time since a given water parcel was last exposed at the water surface. Both observations raise the question of how the bottom water is renewed in such a deep lake. In most lakes, partial or complete turnover of the water column if it occurs - is caused by the seasonal cooling and heating of the water, often supported by wind forcing. Yet, this mechanism is commonly restricted to the upper few hundred meters. In deeper water bodies horizontal density gradients are necessary to induce vertical exchange down to the bottom. Often such gradients are produced by a combination of temperature and salinity gradients. The formation of deep water in the North Atlantic is one of the most prominent examples of this (Dickson et al. 1990). In lakes such gradients are often caused by inflows with differing salt concentrations (e.g. Aeschbach-Hertig et al. 1996a). In deep freshwater lakes in which salinity gradients are small, the non-linearity of the equation of state of water gives rise to two special phenomena. (1) The thermobaric instability is linked to the pressure dependence of the temperature of maximum density (Tmd). It has been proposed by Weiss et al. (1991) as the key process for deep-water formation in Lake Baikal during winter, when the surface water is colder than the bottom water. (2) Cabbeling (McDougall l 987b) results from the non-linear temperature dependence of water density, as a result of which the mixing of water masses of different temperature always leads to an increase in the mean density. The best known example is the so-called thermal bar, where two adjacent surface water masses, one warmer and the other colder than 4 °C, mix to form water of maximum density which then sinks to greater depth (Imboden and Wtiest 1995). Shimaraev et al. (1993) interpret the thermal bar which develops in spring along most of the shoreline of Lake Baikal as the main trigger of deep-water formation. The authors suggest that the thermobaric instability is initiated by the spring thermal bar at the south- eastern shore of the Central Basin. Deep-water renewal is then accomplished by a jet of cold 23 56 53.7 53.6 55 53.5 108 108.4 z 54 ~ ~.a ~ ...i 53 Selenga 52 Lake Baikal 0 50 100 km 51 104 105 106 107 108 109 110 Longitude [0 E] Figure 3.1: Map of Lake Baikal showing isobaths at 400, 700 and 1000 m depth. The lake is divided into Southern, Central and Northern basins: SJ, Cl and NJ mark the deepest points of these basins. White circles show the locations of further sampling stations. The cold bottom boundary layer found in spring 1995 in the Northern Basin is shaded. The Selenga, Upper Angara, and Barguzin are the main inflows, the Angara at the south-western end of the lake the only outflow. The transect between Olkhon Island and the shore of Boldakovo is marked by small black dots. The insets show the subsurface sills at Academician Ridge and Selenga Delta which separate the three basins. Kl to K5 are sampling stations in the Kukui Canyon cut into the Selenga Delta. Sampling stations across Academician Ridge are labelled Al to A4. 24 water which supposedly originates from the hypolimnion on the open-lake side of the thermal bar, However, it remains unclear how the cold surface water is able to cross the TnuJ-line and sink to the bottom. Furthermore, the temperature of the bottom water in the Central Basin is colder than the jet proposed by Shimaraev et al. ( 1993). In fact, a more detailed analysis of the thermal bar leads to the conclusion that it can only explain vertical mixing in the top 200 to 300 m of the water column (Peeters et al. 1996). Since T md decreases with depth (pressure) by about 0.2 K per 100 m, the water at 4 °C which forms at the thermal bar cannot sink very far before it meets with colder ambient water that, at this depth, is denser. Walker and Robert (1995) have developed a numerical model to study the intensity and scale of plumes involved in the formation in deep water. Deep convection is assumed to be triggered by a storm surge which forces cold surface water below its compensation depth (where water temperature is equal to the local TmJ). This phenomenon is simulated by initiating the model calculations with an unstable temperature profile in which the mesothermal temperature maximum (MTM), i.e. the layer in which temperature reaches its maximum and crosses the depth-dependent Trrn1-curve, is located deeper than 500 m. If the initial instability is applied adjacent to a solid boundary, the model simulates vigorous plume motion and the production of a large volume of deep water. Unfortunately, the conditions used for the model calculations, especially the extremely deep MTM, have never been observed in Lake Baikal. Thus, it remains unclear whether the model would succeed in simulating a sinking plume containing water that is really cold enough to explain the observed temperature decrease in the deep water of the Central Basin using more realistic initial conditions. Killworth et al. (1996) use a two-dimensional model to show that wind forcing is a possible trigger for deep convection. By continuously applying a sinusoidal wind stress to the entire lake surface with a random stress added, complete overturning of the water column is achieved. The authors then use an inverse 11/2-dimensional filling-box model to calculate annually averaged fluxes by assuming the observed distribution of tracers to be stationary (oxygen) or to have a stationary relative vertical structure (CFCs). The vertical fluxes calculated from the model decrease with depth. This could result from wind-induced mixing, but is also consistent with the concept of a river plume which, depending on its temperature, sinks to different depths. At a fixed depth the calculated fluxes vary by a factor of up to two between the different basins suggesting that the process of deep-water renewal is different in each basin. There is a basic difficulty linked to the process of thermobaric instability proposed by Weiss et al. (1991) and Killworth et al. (1996). Since water temperature near the bottom of the Central Basin drops to about 3.1 °C, the MTM would have to be pushed downwards by about 350 m to cause water of 3.1 °C to sink from the surface to the lake bottom. Although 25 wind shear cannot be specifically excluded, several hundred CTD profiles taken during different times of the year have only rarely shown thermocline displacements of more than just one fourth of the required distance. For instance, measurements before and after a severe storm in November 1994 with wind velocities above 20 m s-l document a temporary downward displacement of the thermocline by just about 80 m. Since salinity variations are extremely small in Lake Baikal (Falkner et al. 1991 ), the influence of dissolved chemicals on water density has been neglected by all authors. However, the thermal expansion coefficient is very small in the region of the MTM, and even small salinity differences may become dominant for the vertical density structure. Thus, deep- water renewal is likely to occur in regions where water masses of different salinity meet. Such a location is found at Academician Ridge, where fresh water from the Northern Basin mixes with slightly more saline water from the Central Basin. Since some of the inflows have a greater salinity than the lake, the river mouths, e.g. the Selenga Delta, are other areas of potential deep-water formation. Based on a thorough theoretical analysis of stability and neutrally buoyant transport in lakes (Peeters et al. 1996) and on a rather detailed set of data collected during several expeditions to Lake Baikal between 1993 and 1995, we are able to (1) confirm from tritium/helium water ages (to be published elsewhere) the order of magnitude of deep-water exchange given by Weiss et al. (1991), and (2) present a more consistent picture of the processes which cause the exchange of deep water. The latter point will be addressed here. Acknowledgements We thank our colleagues of the Limnologica/ Institute of the Siberian Division of the Russian Academy of Sciences in Irkutsk, especially its director M. Grachev, and the Baikal International Centre of Ecological Research (BICER) for providing ship time and support. Special thanks go to Michael Schurter for his reliable technical and experimental work, to T. Khodzher and Laura Sigg for providing chemical data and to David Llewellyn-Jones for providing satellite images. We also thank N. Granin, A. Zhdanov and the crew of the RV Vereshchagin. This research was made possible by the financial support of the Swiss Federal Institute of Environmental Science and Technology (EAW AG), the Swiss Federal Institute of Technology (ETH) and the Swiss Federal Office for Science and Education (BBW). 26 3.2 Methods The study described here is based on high-resolution CTD profiles taken during several cruises on Lake Baikal, complemented by chemical analyses of water samples. This information is used to calculate one- and two-dimensional distributions of water density, which are then analysed in terms of gravitational instabilities and the potential for the formation of density-induced currents. There are three steps involved in this procedure: ( 1) calculation of "salinity" S from electric conductivity and chemical water analysis; (2) calculation of the salinity effect on water density (the coefficient of haline contraction) for the specific chemical composition of Lake Baikal water; and (3) evaluation of the calculated density field with respect to gravitational flow and mixing. 3.2.1 Calculation of Salinity from Conductivity Salinity Sis defined as the total mass of dissolved solids per unit mass of solute. In principle, a water sample should be analysed for all its major components and its salinity S [g kg-I] calculated as follows: l S=-:EM;·C;, (3. l) Pi where p is the water density [kg m-3], M1 the molar mass [g mol-1] and c; the molar con- centration [mol cm-3] of species i. If the relative composition of the water is constant, Scan be calculated from the electrical conductivity K. However, depending on the chemical composition of the water, each lake has its own specific KIS ratio which can be derived from basic physico-chemical properties of the dissolved ions as described by Wiiest et al. (1996). Firstly, the conductivity measured at temperature T and pressure p, Kif., is normalised to standard pressure: (3.2) where K¥ [µS cm-1) is the conductivity at atmospheric pressure (conventionally defined as p = 0). The specific pressure correction o(T) accounts for the increase in ionic conductivity due to compression. Since no data are available for freshwater, Wiiest et al. (1996) linearly extrapolated the sea-water data of Bradshaw and Schleicher (1965) to zero salinity. Their equation (4d) applied to T= 3.5 °C (a typical value for the deep water of Lake Baikal) yields o(3.S °C) = 1.669· I0-5 dbarl. To check this value we took vertical CTD profiles with the conductivity and temperature sensors wrapped in a plastic bag to prevent water exchange with 27 Table 3.1: Typical ionic concentrations c; and conductivity in Lake Baikal water. A.~ c CJ a Z; 20,i F;d "'20,i e [mmol kg-I] [mS cm-I (eglltl] H {µS cm-I] HCO=j l.0921 b 40.29 0.97 42.69 co~- 0.0005 b 2 63.18 0.93 0.05 c1- 0.0123 68.81 0.98 0.83 sol- 0.0574 2 71.86 0.93 7.71 Na+ 0.1550 44.80 0.97 6.76 K+ 0.0241 66.52 0.98 1.57 Ca2+ 0.4020 2 53.04 0.92 39.42 Mg2+ 0.1260 2 47.00 0.92 10.89 1.: = 109.91 a Concentrations from Falkner et al. (1991) b Calculated from alkalinity= 1.093 molfk.g, pH= 7.2 (Falkner et al. 1991) and T= 3.5 °C c Equivalent conductance A-2Q,; at T = 20 °C and infinite dilution. For details see Wuest et al. (1996) d Reduction coefficient F;. For details see Wuest et al. (1996) e Conductivity of ionic species i at T = 20 °C, lr2Q,i = p · F; A.io.i ·Z; · c; [µSiem], with p = density of water and Z; = ionic charge the surrounding water. Temperature within the bag was approximately constant ( (J "'3.5 ± 0.02 °C). The experiment yielded oexp(3.5 °C) = l.556-10-5 dbar·1. The correspondence between the extrapolated and the experimental value is remarkable. We prefer the latter value since it includes the possible instrument-specific effect of pressure on the conductivity sensor. In fact, it appears that the temperature dependence of o(T) is not relevant for the range encountered in Lake Baikal. Secondly, the reference conductivity at T =20 °C, KSo , is calculated as follows: (3.3) The polynomial fr is determined from the specific ionic composition of Baikal water measured by Falkner et al. (1991) and the temperature dependence of the equivalent conductance of the major ions (Tab. 3.1; see Wuest et al. (1996) for details). For Baikal water: 2 4 2 fr= 0.59911+1.6899· 10- ·T+1.9024·10 · T -1.6495· I0-6 · T3 (3.4) (Tin °C). 28 Finally, the salinity due to the ionic composition of the water, Sc , follows from ICgo by applying (3.5) where Cs is calculated from the equivalent conductance of the major ions. Since the ionic conductance decreases with the ionic strength of the solution, gs increases with salinity, i.e. with electric conductivity. For Baikal water it can be approximated by the polynomial gs=8.4456·10-4 +3.2654·10-7 ·IC20-4.3774·100 -10 ·( IC200 )2 (3.6) The notation Sc was introduced in Eq. (3.5) to distinguish it from the "non-ionic salinity" So. Among all the chemicals that are subsumed under So, in the context of this investigation only Si(OH)4 contributes significantly to the spatial variation of water density. The corresponding salinity is therefore called Ss;. Total salinity Sis then defined by (3.7) 3.2.2 Coefficients of Haline Contraction and Equation of State The coefficients of haline contraction for the ionic species and for silica, Pc and Psi respectively, are defined by: (3.8a) 1 ap Psi=--. (3.8b) p assi These coefficients can be calculated for standard conditions (atmospheric pressure p = 0, T= 25 °C) from the partial volumes of the dissolved solids and the specific chemical composition of the water (W!iest et al. 1996). For Baikal water: Pc(T = 25°C, p = 0) = 0.8107·10-3 (g kg-1)-l; (3.9a) 3 Ps;(T = 25°C, p = 0) = 0.36·10- (g kg-1)-1. (3.9b) To calculate the coefficients for arbitrary T and p we assume that the temperature and pressure dependence of Pc and Ps; are proportional to the corresponding variation in the 29 coefficient of haline contraction for sea-water, f1cM, calculated from the equation of state by Chen and Millero (1986): Pc(T,S, p) = l.074 ·PcM(T,S,p); /3s;(T,S,p) =0.477 · PcM(T,S,p). The full equation of state for Baikal water p(T, Sc, Ss;. p) can now be calculated by integration. Since the functions in the following integrals are nearly constant, they can be approximated by simple multiplication. Thus: For the whole Jimno)ogical range (0 °C:::; T:::; 30 °C; 0 mg kg-I:::; S:::; 600 mg kg-I; 0 bar :::;p ::> 180 bar) the values calculated using Eq. (3.10) and those calculated using the polynomial by Chen and Millero (1986) agree to within 8 significant figures. For the range found in Lake Baikal the values match within 10 significant figures. 3.2.3 Water Density and Stability of the Water Column A profound description of the theoretical concepts underlying the calculation of stability from water density is given by Peeters et al. ( 1996). Equations from this publication are cited as (PE plus equation number). The equation of state can be formulated in terms of in situ values of salinity S, pressure p and either water temperature Tor potential temperature e(PE 7): p(8,S,p) = p(T,S, p), (3.11) where """ denotes a function of e. s and p. Variations in density due to temperature are described by the coefficient of thermal expansion, a(T,S,p) or a(8,S,p) (see Gill 1982 and PE 9, 10). In freshwater at atmospheric pressure, maximum density is reached at 3.98 °C. The temperature of maximum density Tmd• at which both a(T, S, p) and a( 8, S, p) pass through zero and change sign, decreases with increasing pressure at a rate of -0.021 K barl. The effect of pressure on a( 8, S, p) for different temperatures and S =0.1 g kg-1 is shown in Tab. l of Peeters et al. (1996). 30 The stability of a water column is determined by the vertical variation in water density. Usually, local stability (the stability with respect to an infinitesimal vertical displacement of a water parcel) is described by the Brunt-Viiisiilii frequency (Gill 1982, PE 16), which, for the case of Lake Baikal, is supplemented by a term describing the effect of silicic acid (z is positive upwards): (3.12) In order to assess stability over a finite depth interval, oceanographers use "potential density" to account for the effect of adiabatic temperature changes due to decompression. This concept can be misleading if the adiabatic transport of a water parcel occurs in the vicinity of T md• especially if a changes sign, which is likely to occur in a deep, cold, freshwater lake (see Fig. 2 and Tab. I of Peeters et al. (1996)). The "quasi-density" Pqua is introduced as an alternative concept (PE 23, 24). The vertical gradient of Pqua is proportional to the Brunt-Viiisfilii frequency at all depths (while for potential density this is only true at the reference depth). In a stable water column, Pqua must increase monotonically with depth. The difference between the ambient in situ density and the density of a given water parcel which is isentropically moved to different depths (PE 30), (3.13) can be used to evaluate the depth to which the fluid parcel can sink or rise from its initial position if no mixing between the parcel and its environment occurs. Here Ppa( z) = P( epa• Spa> p( z)) is the density of the parcel with constant potential temperature epa and salinity Spa but variable pressure p( z), and p( z) is the in situ density of the ambient water. For the case of a two- or three-dimensional (e, S)-field, McDougall (1984; 1987a) introduces the "neutral surface'', along which infinitesimal displacements are neutrally buoyant. Over finite distances, however, a water parcel moves off the neutral surface. In contrast, the "neutral track" is defined in (PE 32) as the path along which a fluid parcel moves buoyancy-free over a finite distance from its equilibrium position. Since the transport along the track is considered to be isentropic, (}pa and Spa remain constant, i.e. mixing with the surrounding water is excluded. Along the neutral track the density of the parcel is always equal to the local density of the surrounding water (PE 29), and thus the neutral track of a given water parcel is defined by l:lp( z) = 0. Note that every water parcel has its own individual neutral track. 31 3.3 Field Data The data presented here were collected during three cruises on Lake Baikal (18 May - 26 June 1993; 21 October - 14 November 1994; 11 May -4 June 1995) which were organised by the Baikal International Centre for Ecological Research (BICER). A total of 581 CTD casts (198 in 1993; 126 in 1994; 257 in 1995) were taken from aboard the research vessel RV Vereshchagin. Water samples for the analysis of tritium, helium isotopes and other noble gases were collected during all three expeditions. These results will be discussed elsewhere. Temperature, conductivity and pressure (depth) were recorded in situ with an SBE-9 on-line CTD probe from Seabird Electronics. The probe provides a resolution of 0.025 dbar for pressure, 0.0003 K for temperature, 0.01 µS cm-1 for conductivity, and 0.01 mg 1-1 for oxygen. Oxygen concentrations were calibrated using Winkler titration data (accuracy: 0.3%) provided by T. Khodhzer (unpublished data). The probe was also equipped with a transmissometer (Sea Tech Inc., path length 10 cm in 1993; Chelsea Instruments Ltd., path length 25 cm in 1994/95). Light transmission was not calibrated for absolute suspended particle concentration; it is used here only as a qualitative tracer, e.g. for river water. Ionic salinity Sc was determined from electrical conductivity according to the procedure described above. Compared to Sc the precision and spatial resolution of silicic acid concentrations measured by T. Khodzher (unpublished data from 1993) is modest. Thus, a mean Ss; distribution was constructed by assuming a constant value in the surface layer (Weiss et al. 1991, Killworth et al. 1996) determined as the mean value of more than 450 samples. Below 300 m the concentration of Si(OH)4 increases with depth. For each basin the vertical profile is approximated by a second-order polynomial (Tab. 3.2). Table 3.2: Polynomials to approximate the basin-specific mean vertical concentration of silicic acid. Si(OH)4 concentration in the top 300 m is approximated by a constant value for the whole lake. (h =distance from water suiface) h[m] Si(OH)4 =a·h2 + b·h + c [mg 1-l] a b c 0-300 all basins 3.08 Southern Basin 1.81-10-6 -1.08· I0-3 3.24 >300m Central Basin 1.89· I0-6 -1.13·10-3 3.25 Northern Basin 4.02·10-6 -2.41· 10-3 3.44 32 3.3.1 Density Stratification in Lake Baikal Potential temperature (in the following, referred to simply as temperature) and water density were calculated from the CTD data and the vertical distribution of silicic acid. The distribution of silicic acid in the years 1994/95 was assumed to be the same as in 1993. In Fig. 3.2a temperature profiles taken at the deepest part of each basin in spring 1995 are shown. All profiles show inverse stratification with respect to temperature in the top 200 m of the water column. The MTM is located at 180 m in the Northern Basin (temperature at MTM, TMTM = 3.59 °C), at 200 min the Central Basin (TMTM= 3.56 °C) and at 195 min the Southern Basin (TMTM= 3.59 °C). Below the MTM, temperature decreases steadily. In the deepest 20 to 50 m of the Central Basin temperature decreases by 0.05 - 0.1 °C. As shown below such bottom layers were also found in the other basins. Ionic salinity (Sc) is low and varies by less than 1 mg kg-I within the water column and between the basins (Fig. 3.2b). The salinity signal measured in the Central Basin seems to be disturbed by intrusions. The destabilising effect of the slight decrease in Sc with depth in the interior of the Southern and Central Basins is small and compensated for by the gradient of 8 (Fig. 3.2a). Salinity in the Northern Basin is significantly smaller and increases with depth. The steady increase in quasi-density with depth in all three basins (Fig. 3.2d) illustrates that the combined effects of temperature, salinity and silicie acid yield a stable stratification. The therrnocline at about 100 m depth marks the zone of maximum stability, i.e. the zone with the largest gradient of Pqua· This gradient is small near the MTM, indicating that the stratification is weak. A pronounced increase in Pqua at the bottom of the Central and Northern Basins points to stably-stratified boundary layers. The contribution of 8, Sc and Ss; to the local vertical stability (Eq. 3.14) is shown in Fig. 3.3. Stability is dominated by the temperature stratification (a· ae/ az ), except near the MTM, where temperature crosses the Tmd curve and &: becomes zero. The contribution of dissolved ions 10 vertical stability, given by -A asc;az. is mostly small, sometimes even negative, except in the top 300 m and in the bottom boundary layer. The contribution of smcic acid to vertical stability, given by -fis; as8;1az, is positive and of intermediate size. Note that the approximation of Si(OH)4 by an average profile (see Tab. 3.2) makes this term rather smooth. 33 0 . ,. 200 N '·"'· 400 .. $ ] 600 N\ 800 l -=c.. : Q 1000 1200 / 1400 ,'a> c b) 1600 1.5 2 2.5 3 3.5 94.5 95.0 95.5 !H°CJ S [mgkg.11 c 0 200 400 l I 600 \ s ] I 800 :9c.. l Q 1000 \ \ 1200 d)c 1400 c c) 1600 10 11 12 13 14 0.055 0.06 0.065 0 [mg1·11 p - 1000 [kg m·3J 2 qua Figure 3.2: (a) Potential temperature e, (b) ionic salinity Sc, (c) dissolved oxygen con- centration [Oz], and (d) quasi density Pqua at the deepest point of the Southern Basin (dashed line), the Central Basin (solid line) and the Northern Basin (dotted line) measured between May 28 and June I, 1995. The straight dashed line shown in (a) is the temperature of maximum density, Tmd· 34 0 r································· 200 ...... 400 '··:.:;:. 600 \ ] Figure 3.3: Impact of potential temperature \ (I ., 'l. 800 9 (dotted line), ionic salinity Sc (solid line), '5a. \ ·Ps; dS1 dz_ ••••• :: \I) ~ and silicic acid Ssi (dashed line) on local 0 1000 \ ...... • \ !~·.,. vertical stability calculated from Eq. (3.14) 1200 ...... for the profile in the Central Basin shown in ·<·a d0/dz Fig. 3.2. The figure is drawn such that 1400 ...... ::, .. positive values correspond to positiv§ 41...... stability. Note that the difference between f3 1600 and /3 is extremely small. ·1 0 1 2 3 4 5 6 10·9 [m.1] 3.3.2 Signals in the Bottom Water In all three basins distinct temperature and salinity signals in the bottom layer are observed, though at different times and of different kind (Fig. 3.4). These signals may help to identify the processes responsible for deep-water renewal and the origin of the water which is mixed to the bottom of the lake. Temperature and salinity changes in both directions were observed in the bottom layer of the Southern Basin (Figs. 3.4a, b). In contrast to the other basins, the signals in the Southern Basin are restricted to the lowermost 3 5 m of the water column. A cooling event occurred between June l and June 26, 1993 which led to a temperature drop of 0.0 l K and a slight increase in the oxygen concentration, but left salinity virtually unchanged. In contrast, the first profiles taken during the cruise of spring 1995 revealed a distinct increase in temperature and salinity as well as a slight decrease in oxygen concentration and light trans- mission in the bottom layer. The temperature increase from 3.375 °C to 3.385 °C causes a density change of --0.35 g m-3 (a"' 35· IQ-6 K-1), which is partially balanced by a change in salinity from 95 to 95.2 mg kg-I, causing a density change of +0.16 g m-3 (/Jc= 0.81· l0-6 (mg/kg)-1 ). The background field of Si(OH)4 causes a small density increase of 0.03 g m-3. Assuming a stable stratification, the remaining difference must be balanced by an additional increase in the concentration of silicic acid and/or suspended particles. In fact, light transmission drops from 79.6% to 78.9% in the lowermost 20 m, indicating that the bottom layer is enriched with particles. 1300 ! l a) f t b) 1350 ·~ i \ ] lJune 193 \ ! lJune '95 i. Southern Basin ~~t ... ..s:::... j/ !. 1:1.. ·:i 'I: Q.. ~ 'I. 1400 26June 193}. r ~ )i -..,'\'" :"' : ,.:...... i .:"• ..;- - - -, 15May '95 1450 3.36 3.37 3.38 94.9 95.0 95.1 95.2 1 0 [°C] Sc [mg kg" ] 1300 •I c) /' d) :, i.:l 1400 I j ] f I fl:; ; i Central Basin 22May'95 : ( \ ! 1:1.. f1 5 Nov. '94 =.. 1500 ~ \\ Q ). \., ' : \' 1)~ 31 May '93 ....,, ...... ~· · 1600 ...... ,..,-/ ...... < , ... 3.1 3.15 3.2 3.25 95.0 95.2 95.4 1 (j [°C] s. [mgkg" ] 700 750 e) f) ] 800 25May'93 =..1:1.. Q 850 1:!.~~=-~.:L-·· .. ,..... 900 24 May'95 3.3 3.35 3.4 3.45 3.S 94.6 94.7 94.8 1 0[°CI s, [mg kg" ] Figure 3.4: The evolution of bottom boundary layers in the three major basins of Lake Baikal traced by potential temperature 9 and ionic salinity Sc- (a, b) In the Southern Basin, formation of a cold bottom layer with virtually unchanged salinity was observed in June 1993, while in May 15, 1995 a warm, saline layer was formed which had disappeared by June 1. (c, d) ln the Central Basin a cold and saline bottom layer was found at all times. (e, /) ln the Northern Basin bottom layers were cold while salinity changed in either direction. 36 In the Central Basin (Figs. 3.4c, d), the bottom layer always consists of cold, higher salinity water enriched in oxygen and has a lower light transmission than the bulk water above. The signals were rather similar in spring 1993 and 1995, but less pronounced in fall 1994. Both increases and decreases in temperature and salinity were observed in the bottom layer of the Northern Basin (Figs. 3.4e, f). A dramatic cooling of the bottom water was registered between May 25 and June 19, 1993 (25 days), when temperature in the lowermost 20 m dropped by more than O. l K, salinity decreased by about 0.2 mg kg-1, and the oxygen concentration increased by about 0.5 mg 1-1. On November 7, 1994, the measurements revealed a pronounced bottom layer with a homogeneous salinity excess, yet no signal in either temperature or oxygen was seen. On May 24, 1995, another cold bottom layer was detected. 3.4 Discussion The temperature decrease with depth below the MTM implies that large scale convective mixing - and not merely diffusion must be involved in the formation of deep water. The supply of cold water must be sufficient to balance the turbulent vertical flux of heat as well as geothermal heating. Nowhere below the MTM is the water cold enough to be able to cool the deep water to the extent observed in the Central Basin (Fig. 3.2a). Thus the cold signal measured in the bottom layer of this basin must be generated either by cold water from above the MTM during winter, or by cold inflows. Since the adjacent basins are warmer, the cold bottom layers in the Southern and Northern Basins could, among other possibilities, originate from the Central Basin. Furthermore, the warm signals found in the Northern Basin - and also occasionally in the Southern Basin - could be produced by geothermal heating (Golubev et al. 1993), by sub-aquatic hydrothermal springs (Kipfer et al. 1993, Kipfer et al. 1996) or by entrainment of warmer water from the surface. A detailed analysis of the depth and the temperature of the MTM reveals two potential regions of deep-water formation: (1) the Selenga Delta between the Southern and the Central Basin; and (2) Academician Ridge, which separates the Central from the Northern Basin. Fig. 3.5 shows the two-dimensional (9, S)-distribution obtained from a transect along the thalweg (the deepest possible path along the major axis) from the northern to the southern end of the lake recorded between June 17 and June 26, 1993. There are several notable features. In the region of both sills, the bottom water is significantly colder than in the open lake at the same depth. Water near the Selenga Delta has a larger salinity than water in the Southern and the Central Basin. Cold and saline water seems to plunge to the deep part of the lake on either 37 0 600 ] :S Q.. ~ 1000 1200 0 100 200 300 400 500 600 700 Relative distance [km] Figure 3.5: Two-dimensional isopleths of (a) potential temperature (J and (b) ionic salinity Sc measured along the tha/weg from north (right) to south (left) on June 17 - 26, 1993. SD = Selenga Delta, AC= Academician Ridge. 38 side of the delta. At Academician Ridge the salinity difference between the Central and the Northern Basin causes the isohalines to become vertical. Saline water seems to sink from the Central Basin to the deep part of the Northern Basin. In the following discussion the situations at both locations are analysed in more detail. 3.4.1 Selenga River The Selenga is the major inflow to Lake Baikal. Its rather large delta is located at the sill between the Southern and Central Basins. Kukui Canyon is cut into the delta and leads down to the deep part of the Central Basin (see inset, Fig. 3.1). CTD profiles measured in the canyon on May 15, 1995 (Fig. 3.6), just after the ice had disappeared from this part of the 0 200 :..~.:~:;r. l .. ····'f ,,.. I t I ;· \ ] 400 1 l ..c: K3 ,,._f L, K3 K3 .... 1 ... ,. .,,,. ,,,. K4! "'...... ,; : K4 ~ 600 ,-& !;"' ~ l 0 I .. .. I .. --1' I ~ I 800 Kf" / I Cl! a) ·Cl "'.Cl b) c) I I I 1000 3.1 3.2 3.3 3.4 3.5 3.6 95 100 105 50 60 70 80 1 9 [°C] S, [mg kg" ] Light transmission [%] Figure 3.6: (a) Potential temperature 8, (b) ionic salinity Sc, and (c) light transmission in Kukui Canyon measured on May 15, 1995. See Fig. 3.1 for positions of stations. The profiles show cold water from Selenga with high salinity and turbidity (low light transmission) flowing along the bottom of the canyon. The bottom layer weakens with growing distance from the river mouth. At position K2 (solid line), located about l 2 km from the river mouth, the bottom water is colder than Tmd· At K3 (dashed line), 7 km further along the canyon, the water is warmer than the local Tmd· At K4 (thick dotted line), 12 km north of K3, only a weak temperature signal remains. A profile measured at the deepest part of the Central Basin (Cl) on May 16, 1995 is shown for comparison (alternating dashes and dots). 39 lake, show water at the bottom of the canyon with high salinity, increased turbidity (low light transmission) and enhanced oxygen concentration. Water with similar characteristics was found at the southern slope of the delta. Measurements at position S2 (see Fig. 3 .1) on May 31, 1995 show a layer of intruding water at 790 m depth with increased salinity and turbidity (Fig. 3.7). Similar structures were found on June 25, 1993. In the Southern Basin the intruding water is slightly warmer than the ambient water (Fig. 3.7), while the bottom water in Kukui Canyon is significantly colder than the ambient water (Fig. 3.6). In fact, the temperature measured at the bottom of position K2 (T = 3.07 °C at 420 m) is the lowest ever registered below 300 m during any of our several cruises. High salinity and low light transmission are indicative of water from the Selenga. Both ionic salinity and concentrations of silicic acid are much higher in the river (Sc = 126 mg kg-1 (Votintsev, 1993); Ssi = 12.5 mg J-1 (L. Sigg, unpublished data)) than in the lake (Sc = 95.3 mg kg-I; Ss; 4.2mg1-1 (L. Sigg, unpublished data)). Light transmission [%] 91.3 91.5 91.7 91.9 ··-·········.... ,_ 200 .., ..... :§ 400 ': 600 i0 800 a) b) c) 1000~...... ~~...... 3.3 3.4 3.5 3.6 95.0 95.5 78.7 78.9 79.1 79.3 1 0 [°CJ S [mg kg" ) Light transmission[%] c Figure 3.7: (a) Potential temperature 8, (b) ionic salinity Sc, and (c) light transmission measured at the southern slope of the Selenga Delta (position S2) on June 25, 1993 (solid line; upper axis in (c)} and on May 31, 1995 (dotted line; lower axis in ( c)). 40 The ionic salinity can be used to calculate the volume fraction, T/mix, of Selenga water in the river plume flowing along the canyon: (3.14) Sc,I is the mean ionic salinity of the lake water, Sc, r the ionic salinity of the river water and Sc, P the ionic salinity measured in the river plume. At position K2 (Sc. P = 107 mg kg-1), the fraction of Selenga water is T/mix = 0.38, and at K3 (Sc, p = 100 mg kg-1) it has dropped to T/mix = 0.15. At position K4, the signal has become too weak to be able to evaluate Eq. (3.14). Unfortunately, no corresponding data for Si(OH)4 exist. Thus, the above mixing ratios cannot be confirmed by an independent tracer. Since the temperature variations in the Selenga are too fast and not continuously recorded, it is not possible to check the mixing ratios by applying Eq. (3.14) to water temperature. However, the T/mix values can be combined with the water temperatures measured in the plume (6p) and in the lake (61) to estimate the temperature of the inflowing river (6,). If 61 is approximated by the ambient water temperature at plume depth, the calculated river temperature becomes 6, "'2. 36 °C with the values from K2 and 6, "'2.18 °C with the values from K3. Those temperatures drop if a larger value is chosen for 61 (the river plume is flowing through the MTM where the ambient water is warmer!). In the middle of May 1995 surface temperatures were about 2 °C in the delta region and I °C at the centre of the Central Basin. Measurements by infrared satellites (David Llewellyn-Jones, unpublished data) confirm these values. By comparing the density of a water parcel at variable depth z, Ppa(6pa•Sc,pa•SSi,pa•Z), with the in situ density, the depth can be calculated to which the parcel would sink by free convection if moved isentropically (see Eq. 3.13). Fig. 3.8 shows t::..p( z) calculated for two types of test parcels moved either to position Cl (Central Basin, dotted lines) or to S 1 (Southern Basin, dashed lines). The first test parcel (bold lines) is taken from the bottom of Kukui Canyon at position K2 ( 6pa = 3.07 °C, Sc.pa = 107 mg kg-l). The second parcel has the same salinity but is slightly warmer ( 6pa 3.5 °C, thin lines). For both cases Ss;,pa was approximated by the mean background Si(OH)4 concentration in the respective basin at 420 m. The difference in Si(OH)4 concentrations between river and lake was neglected for calculating t::..p( z) and the sinking depths are therefore slightly underestimated. According to Fig. 3.8, water from Kukui Canyon has the potential i.e., if entrainment is negligible of reaching the bottom of both the Central and Southern Basins. If water temperature is rising - e.g., in late spring when the Selenga becomes warmer - the plume would no longer reach the bottom of the basins. This process starts earlier in the Central than in the Southern Basin. 41 Figure 3.8: Difference between in situ water 0 --....,;""Y density at position Cl (Central Basin, dotted 200 lines) or SJ (Southern Basin, dashed lines), y respectively, and the density of a test parcel, 400 Hf" Ppa(8pa,Sc,pa•Z). moved isentropically to 1! ~.. these profiles (see Eq. 3.15). Bold lines: Test 600 ! If i\ parcel from the bottom of Kukui Canyon at I i \\ position K2 ((}pa 3.07 °C, Sc,pa 107 mg £i 800 I : \ = I ...... kg-1 ). Thin lines: as before, but water \ Ci.."' 1000 I temperature raised to 8pa = 3.5 "C. Note that I i \ if entrainment is neglected the parcel sinks 1200 I i I ! through the water column as long as !J.p(z) 1400 I i remains negative. Thus, for both temperatures \ the test parcel would reach the bottom of the \ 1600 .. Southern Basin. Jn the Central Basin it would -0.02 -0.01 0 0.01 be stopped at an intermediate depth if (}pa 3 were 3.5 °C. !J.~ [kg/m ] Two-dimensional isopleths of (}and Sc, shown in Fig. 3.9 for the conditions met in Kukui Canyon on May 15, 1995, confirm the conclusions drawn from Fig. 3.8. Selenga water flowing through the canyon is traced by its lower temperature and higher salinity. The neutral tracks shown in Fig. 3.9a (bold lines) suggest that the bottom water from position Kl and from position K2 moves along the canyon to the deep part of the basin. The neutral track gives the direction of flow of a parcel that does not mix with ambient water. As shown above, mixing by entrainment is significant. Alternatively, the neutral surfaces (small sticks in Fig. 3.9a) represent the direction of flow of a water parcel which is continually being mixed with its environment. Since neutral tracks and neutral surfaces are fairly parallel to each other, the two complementary concepts define the rather narrow depth range within which Selenga water is expected to flow to the deep layers of the basins. In spring 1993, when measurements started somewhat later in the year, the Selenga plume could be traced in the canyon to a depth of about 700 m. On May 23, 1993, a saline and slightly warmer lens of water could be identified 70 km north of the Selenga Delta at 1000 m depth in a section extending from the eastern shore of Boldakovo to the western shore of Olkhon Island (Fig. 3.10). The lens was also detected three weeks later, on 12 June. Based on observations made at the water surface (see below) we conclude that this water originates from the Selenga. According to the salinity signal, the fraction of Selenga water within the lens would be 11mix =0. 023. If flowing at a mean speed of 3 cm s-1, it would take about 4 weeks for the river water to reach Boldakovo. Note that due to entrainment of warmer water from the MTM and due to the positive temperature gradient in the deep water of 42 200 400 ! ':flc.. 600 QI Q 800 1000 1200 0 200 400 600 800 1000 1200 5 10 15 20 25 30 35 40 45 Distance from inlet [km] Figure 3.9: Two-dimensional isopleths of (a) potential temperature (fond (b) ionic salinity Sc measured along Kukui Canyon on May 15. 1995. Selenga water flowing through the canyon is traced by its lower temperature and larger salinity. The mesothermal temperature maximum (MTM) is poorly developed near the river mouth. The neutral track (Fig. 3.9a: bold line) of a water parcel from the bottom at position Kl shows the potential of riverine water to penetrate through the MTM, although initially it does not flow along the bottom. The neutral track which starts at the bottom of position K2 follows the canyon. The neutral surfaces (small sticks). representing the direction of flow of a water parcel which is continuously mixed with its environment, are fairly parallel to the neutral tracks. Note that for the calculation of the neutral tracks and neutral surfaces the influence of silica was included. 43 the basin, the lens is now wanner than the ambient water, in contrast to the observations made in the Kukui Canyon where the river plume is always colder. In early June, the spring thermal bar (Shimaraev et al. 1993, Weiss et al. 1991) develops at Boldakovo and at other places along the eastern shore. At Boldakovo, the creation and movement of the thermal bar is controlled by the bottom topography and the northerly flow of warm Selenga water along the shore (Shimaraev et al. 1993). Chemical analysis of water samples taken from either side of the thermal bar and from the Selenga (L. Sigg, unpublished data) are applied to Eq. (3.14) to calculate the mixing ratio for the surface water trapped between the thermal bar and the shore (Tab. 3.3). Although the numbers for the different chemical components are not identical (how could they be, given all the uncertainties and the rather coarse resolution, e.g. for Sr?), they nonetheless give a fairly consistent picture, according to which between 20% and 50% (mean= [36 ± 12]%) of the water along the shore originates from the Selenga. The ratio is significantly larger than that determined for the water lens detected at 1000 m depth (Fig. 3.10). This indicates that entrainment of ambient lake water is greatly suppressed if the river water flows at the surface, trapped behind the thermal bar. Despite its high salinity, the water on the shore-side of the thermal bar is too warm (T> Tmd) to sink to the bottom of the Central Basin. This was shown by Peeters et al. (1996) by using neutral tracks and neutral surfaces calculated from two-dimensional isopleths of fJ and S across the thermal bar at Boldakovo. Table 3.3: Relative volume fraction Tl mix of Selenga water in the water trapped between the thermal bar and the eastern shore of the Central Basin calculated from selected chemical data. Cc. Cs and CrB are concentrations in the open surface water of the Central Basin, in the Selenga River and in the water trapped behind the thermal bar at Boldakovo. Samples taken in June 1993 (L. Sigg, unpublished data). Alkalinity (mmol 1-ti l.065 1.435 1.205 0.38 Sr (µmo! 1-l) L5 1.975 1.6 0.21 soi- (mmol J-1) 0.056 0.083 0.065 0.35 c1- (mmol J-1) 0.0145 0.0375 0.026 0.50 i'imix =0.36±0.12 a 11mix CTB -Cc; volume fraction of Selenga water trapped behind thermal bar (Eq. 3.14). Cs-Cc 44 The observations presented in Figs. 3. 7 to 3 .11 can now be combined to yield a joint scenario of the likely sequence of events related to the Selenga. In spring, when, due to snow melting, the discharge of the Selenga increases from its typical winter mean value of about 100 m3 s-1 to a monthly mean of about 600 m3 s-I in April and 1700 m3 s-I in May (Shimaraev et al. 1994), the water in the river plume is cold and saline enough to flow along the lake floor to the bottom of the Central Basin and possibly to the bottom of the Southern Basin. On its way it entrains significant quantities of lake water from different depths. This is reflected in the mixing ratio, which decreases from T/mlt = 0.38 at position K2 to r/mlt =0.15 at position K3. Beyond K3, direct information on entrainment is lacking. However, other data can help to obtain a rough estimate of the "final'' mixing ratio. Firstly, for the lens at Boldakovo (Fig. 3.10) a mixing ratio of T/mix = 0.023 was determined. Secondly, by assuming that the salinity increase of approximately 0.4 mg kg-I measured at the bottom of the Central Basin (Figs. 3.2b and 3.4d) is caused by Selenga water, one gets rfmix"' 0.013. Both values point to a "final" mixing ratio lying between 1% and 3%. In Tab 3.4 absolute flow rates in the plume are summarised for a mean discharge rate of 1700 m3 s-1 (typical value in May). Table 3.4: Entrainment of ambient lake water into the Selenga River plume on its way to the deepest part of the Central Basin. mixing ratio flow entrainment CFC-12 b 1Jmix [ • l [m3 s-1] [m3 s-l] (emol kg-I] Selenga (0 m) 1700 a 2.7 c 0 400m 2800 2.7 K2 (400m) 0.38 4500 400-600 m 6800 2.2 K3 (600 m) 0.15 11300 600-1500m 73700 1.2 Cl [> 1500 ml 0.02 85500 1.32 [1.35 dl • Mean discharge of the Selenga in May (Shimaraev et al. 1994) b Volume-weighted CFC-12 concentrations in the corresponding depth zone determined from the vertical profile measured by Weiss et al. (1991) c Assumed to be equal to the concentration at the lake surface d Calculated from the entrainment rates and the CFC-12 concentrations measured in the upper part of the water column 45 Boldakovo Olkhon 0 ______...... •• 3.0.··· 3.5 200 400 ] 600 c.. -= 0 96.0 - 95.5 •• 95.2 ...... 95.1 200 ' ' .... 95.2······...... ' ' .-· .. 400 ] 600 ...... --...... •. 95,2 ••• ------.... c.. -= 1200 5 10 15 20 25 30 35 40 45 50 Distance from Boldakovo [km] Figure 3.10: Two-dimensional isopleths of (a) potential temperature 8 and (b) ionic salinity Sc along a transect between Boldakovo and Olkhon (see Fig. 3.1) measured on May 23, 1993. The isopleths show a lens of warm, saline water 12 km off the shore of Boldakovo at about 1000 m depth. Based on salinity, the water within the lens contains about 2% Selenga water. 46 One way to check this scenario is to quantify its implications for the vertical transport of a conservative chemical such as CFC-12 (measured by Weiss et al. (1991)). The estimated flow rates in the plume and the measured volume-weighted mean concentrations of CFC-12 in the water column yield 1.35 pmol kg-I for the concentration below 1500 m, which is in good agreement with the observed mean value of 1.32 pmol kg-I (Tab. 3.4). Based on the discharge rates and water temperatures of the Selenga we deduce that during April and about half of May, and during October, (1) discharge is significant and (2) the water has the potential to flow to the bottom of the lake. Total discharge during these two periods is about 5 kffi3. According to Votintsev (1993), about half of the Selenga water flows into the Central Basin, while the other half is diverted to the south. Assuming a "final" entrainment factor of 50 (Tab. 3.4), the water volume brought by the Selenga to the Central Basin (2.5 km3) increases to about 125 km3 while flowing to the deepest part of the basin. This corresponds to the volume of the basin below 1500 m. A similar behaviour is to be expected for that part of the Selenga which flows into the Southern Basin. According to Fig. 3.8 the potential of the riverine water to sink here may be even slightly larger than in the Central Basin. Measurements by infrared satellites in late May (David Llewellyn-Jones, unpublished data) show that Selenga water also flows southwards along the eastern coast. The intrusions detected in the CTD profiles at the southern slope of the Selenga Delta (Fig. 3.7) can be interpreted as a mixture of water from the Selenga and from the lake, which sinks to the deeper parts of the basin. However, in contrast to the Central Basin, so far we have not been able to directly detect a plunging plume here. 3.4.2 Academician Ridge At Academician Ridge, water masses with different (}and S characteristics meet (Fig. 3.11 ). Below 50 m, temperature is generally lower in the Central Basin than in the Northern Basin. The MTM is warmer and shallower in the Northern Basin (TMTM =3.59 °C at about 160 m depth) than in the Central Basin (TMTM =3.55 °C, 220 m). Below the MTM the horizontal temperature difference across the ridge increases to about 0.1 K. Salinity is about l mg kg-1 greater in the Central Basin than in the Northern Basin, since this basin receives the "saline" water from the Selenga, while the Northern Basin is chemically diluted by water from the Upper Angara, which has a salinity of Sc= 81 mg kg-I (Votintsev, 1993). The horizontal salinity gradient between these basins results in vertical isohalines near the ridge. 47 Central Basin Northern Basin 0 1.0 ~_.: 100 --- 200 300 ·3.55 ------] 400 3.52 -:S 500 ------Qia. Q 600 3.4 700 800 900 1000 0 -----~! __ 100 200 300 94.5 ! 400 500 a. =Qi Q 600 700 800 900 1000 0 50 100 150 200 250 Relative distance lkm] Figure 3.11: Two-dimensional isopleths of (a) potential temperature e and (b) ionic salinity Sc along a transect across Academician Ridge from the Central (left-hand side) to the Northern Basin (right-hand side) measured on May 22 - 24, 1995. The 3.42 °C isotherm outlines the cold bottom boundary-layer shown as the shaded area in Fig. 3.1. 48 A distinct temperature decreases by 0.03 - 0.05 K was observed in the lowest 10 to 20 m of the Northern Basin between May 23 and May 28, 1995. The cold bottom layer extends over a distance of about 200 km from Academician Ridge to 55 ° N (shaded area in Fig. 3.1). It covers an area of approximately 4000 km2 and has a volume of 40 80 km3. The cold bottom water is enriched with oxygen, which indicates that it must originate from above the MTM. A pronounced temperature decrease was also detected in spring 1993 (Fig. 3.4e). However, at this time the cold temperature signal was only found in the deepest part of the basin. All attempts to find the cold bottom layer (characterised by()< 3.42 °C isotherm in Fig. 3.11) north of 55 ° N failed. There is no indication that it is established by water from the Upper Angara, since the latter would produce a negative salinity signal. Yet, the 3.42 °C isotherm extends southwards and even rises to Academician Ridge. The temperature at the sill crest (depth 300 m) is equal to the temperature of the cold bottom layer in the Northern Basin. The isotherms suggest that water with a temperature of 3 .40 °C flows down on both sides of the ridge to the deep part of the lake. Whereas in the Central Basin the cold water reaches only intermediate depths of about 600 m, in the Northern Basin it sinks all the way to the bottom. Thus, Academician Ridge is a source of the cold deep water found in the Northern Basin. Measurements along a short transect recorded on May 23, 1995 give a closer view of the stratification at the ridge (Fig. 3.12). The water column on the sill (position A2) is influenced by water from both the Central and Northern Basins. The isohalines and isothe1TIJS reveal a complex structure of interleaving layers of different temperature and salinity. At the ridge, a lens of cold water (T = 3.475 °C) with a salinity of 95 mg kg-I causes the MTM (bounded by the 3.52 °C isotherm) to sink to 270 m depth. The neutral tracks of different water parcels across Academician Ridge are shown in Fig. 3.13. The track of a parcel from position Al at 126 m depth (0) follows the 3.475 °C isotherm to position A2, indicating that its direction is determined mainly by temperature (Peeters et al. 1996). At position A2, the track sinks to the lens of cold water at 160 m bounded by the 3.475 °C isotherm. It then continues along the 94.6 mg kg-I isohaline, suggesting that close to the MTM its direction is dominated by salinity. The track then becomes unstable, i.e. it turns south and then north again. Thus, a parcel moving along this track drops from 230 m through the MTM to 330 m depth (see Fig. 6 of Peeters et al. 1996). The track continues to sink to about 370 m depth at position A3 and finally follows the isotheffilS to the deep part of the Northern Basin. 49 Al A2 A4 0 100 -, I I - -3.475~ - ' I '.... _, 200 3.55 -- ____ 300 ..... ------3.475 - 400 500 600 700 800 """"--...... --- ...... --- ...... --- ...... --..-..._- ...... 0 100 200 .. •94,5..__ ' 'I 300 I ' 400 500 600 700 5 10 15 20 25 30 Relative distance [km] Figure 3.12: Two-dimensional isopleths of (a) potential temperature (J and (b) ionic salinity Sc along a short transect across the Academician Ridge measured on May 23, 1995. Isotherms and isohalines show interleaving water masses from both sides of the ridge. 50 The neutral track of a parcel from 140 m depth at position Al (e) first takes a similar course. Beyond A2, the track becomes unstable and drops from 220 m through the MfM to 350 m depth. Below the MfM the neutral track continues along the bottom to the deep part of the Northern Basin. A third neutral track starting at 200 mat position Al (®) also sinks through the MfM to the deep part of the Northern Basin. Yet, due to the larger initial temperature, the track does not sink as deep as the tracks of the other parcels. An additional track illustrates the path along which surface water from the Northern Basin can move buoyancy-free across the ridge to the Central Basin. It leads from 100 m depth at position A3 (<))to 126 mat position Al. In contrast to the track of the upper parcel from position Al, it does not drop to the cold water lens, but remains close to the 3.5 °C isotherm at position A2. Al A2 A3 A4 0 Jllf~///\\I II ' \ I 5 10 15 20 25 30 Relative distance [km] Figure 3.13: Neutral tracks (thick solid lines) and neutral surfaces (small lines) across the Academician Ridge calculated for the situation encountered on May 23, 1995. The symbols mark the initial position of four different neutral tracks. Some isotherms and isohalines from Fig. 3.12 are included. Note that these results should not be over-interpreted. In fact, the bottom topography and the horizontal resolution of the measurements determine the course of the neutral tracks to a great extent. For instance, imagine that position A2 were located at the slope rather than at the top of the sill. Then the neutral track from 140 m would hit the bottom south of the sill 51 and sink into the Central Basin. Thus, the above discussion is meant to demonstrate the very special situation existing at the sill and the potential for the flow of water from the sill into both basins. Furthermore, one should remember that the concept of neutral tracks is based on the assumption that the parcel does not mix with the surrounding water. In contrast to this, the neutral surfaces describe the displacement of water which always mixes totally with its environment. At Academician Ridge both neutral tracks and neutral surfaces are nearly parallel. Both concepts suggest that water from above the MTM of the Central Basin is capable of sinking into the deep part of the Northern Basin. From the data it appears that deep-water formation takes place over a wide area of Academician Ridge. However, based on the available data a precise quantification of the overflow is not possible. Six days later, on May 29, 1995, the situation encountered at Academician Ridge was practically the same as the one shown in Figs. 3.12 and 3.13. Neutral tracks calculated for water parcels from between 125 m and 200 m depth at position Al sink to the deep part of the Northern Basin. In 1993, when measurements started later in the year, surface temperatures at the ridge were already 3.4- 3.6 °C. On June 20, 1993, neutral tracks calculated for water parcels at the Central Basin side of the sill between 150 and 200 m depth still sink below the MTM of the Northern Basin but stay at 450 - 550 m depth. There is no indication of a similar process taking place at the sill between the Southern and Central Basins. The salinity difference between the basins is too small to cause a significant density difference. Therefore, water from one side of the sill is not able to penetrate through the thermocline and the MTM on the other side. Furthermore, horizontal stratification of the water column in the delta region is dominated by water from the Selenga. 3.5 Conclusions The physics of Lake Baikal is unique in several respects. The lake is the deepest body of fresh water on Earth. Except for the top I 00 m, temperatures are always close to the in situ temperature of maximum density, and thus the thermal expansivity of the water is small. Salinity is small and rather homogeneous, despite the fact that it differs by a factor of two between the major inflows. Finally, the vertical gradient of Si(OH)4 contributes significantly to the vertical stability of the water column. As a result, density variations in the lake are extremely small and three-dimensional. They are the result of a delicate interplay of physical and chemical factors, although mostly - but not always - the variation in e is dominant. To our knowledge, only in Crater Lake (Oregon) are similar conditions found (McManus et al. 1993). 52 Based on the concept of "neutral tracks" developed by Peeters et al. (1996) and an extensive set of CTD profiles, we were able to identify the conditions and mechanisms of deep-water exchange. This process takes place where water masses with different properties meet horizontally. Although it would be premature to say that all such locations and time periods were found, we think that two important sites have been identified: (l) the Selenga Delta and Kukui Canyon leading from the delta into the Central Basin; and (2) Academician Ridge, separating the Central and Northern Basins. Important characteristics of the Selenga are its relatively high salinity and its large annual variation in discharge rate and water temperature. The latter leaves a short time slot in April and early May when the river water is not yet too warm and discharge is already large enough to form a plume which sinks to the deepest part of the Central Basin. Semi- quantitative estimates of the total mass flow per year confirm that the process described is in good agreement with observations of tracers such as chlorofluorocarbons (CFC). A similar process is likely to occur in the Southern Basin but bas not been observed so far. Deep-water formation at Academician Ridge is triggered by the small salinity difference between the Central and Northern Basins which, in tum, reflects the salinity difference in the major inflows to these basins. Here the fluxes are more difficult to quantify as long as measurements with higher spatial and temporal resolution are missing. In Lake Baikal, annual exchange of water and salt is very small compared to the large volume of water. Therefore, spectacular changes in the physical and chemical characteristics of the lake are not to be expected. It is also very difficult to detect small residual fluxes in the total salt balance on a year-to-year basis. At present, it is not clear whether the salinity budget of the lake is balanced. Given the extreme sensitivity of deep-water mixing with respect to the salinity gradients in the lake, it is easy to see that vertical mixing may be strongly influenced by possible Jong-term changes in the salinity of Lake Baikal. Once confirmed by measurements in the lake, such changes are difficult to stop and reverse within a short time. Therefore, in order to protect this unique body of water, methods have to be developed to anticipate possible harmful changes in Lake Baikal. With respect to vertical mixing and oxygen distribution two things are important: (1) continuous monitoring of Baikal's major inflows to detect possible changes in the salt balance of the lake; and (2) collection of more information on the physico-chemical properties of the water body. Unless a policy is adapted to monitor the lake in this way, developments which may lead to irreversible changes in this very special ecosystem may be identified too late. 4. Thermal Bar 4.1 Introduction In large temperate lakes where water temperature crosses the temperature of maximum density twice yearly, thermal bars develop in spring and late autumn. Shallow temperate lakes are dimictic, i.e. they tum over twice yearly. In deep temperate lakes only the surface layer undergoes this biannual change. Since a thermal bar can only exist when two water masses, one colder, one warmer than Tmd• meet at the same depth, thermal bars obviously cannot occur below the MfM, where water temperatures always exceed than Tmd. In early spring, shortly after break-up, the water column above the MfM is inversely stratified with respect to temperature. As air temperature increases in spring, the surface waters heat convectively. In the deep central part of the lake, convective mixing is limited in depth by the MTM, whereas in the shallow near-shore regions the depth limit is given by the bottom topography. Due to this variable depth limit and due to differential heating and wind sheltering, temperatures in the shallow areas increase faster than in the deep mid-lake regions. As soon as temperature in the shallow inshore region exceeds 4 °C, a stable summer stratification starts to develop. Between the warm region near the shore and the mid-lake region where convection still persists, there is a zone where water the temperature Tmd is attained, and water masses sink. This zone is referred to as a thermal bar. With increasing water temperatures, summer stratification intensifies and the thermal bar migrates slowly towards the central part of the lake. Finally, when temperatures in the entire lake exceed T md• the thermal bar disappears and summer stratification develops. In autumn, the thermal bar process develops in the opposite direction. Convective cooling causes the stable stratification to diminish and a convective layer to establish. As autumn cooling progresses, near-shore waters reach temperatures below Tmd before the water masses in the central part of the lake. A thermal bar develops separating the inshore region where winter stratification starts to develop from the mid-lake region where autumn 54 convection persists. As autumn cooling continues, the thermal bar migrates towards the central par of the lake. Further cooling will cause the stratification to intensify and finally the lake starts to freeze. The thermal bar was first observed in Lake Geneva during the late 19th century by F. A. Forel (1880) who referred to it as a "barre thermique". More recent observations of the evolution of the thermal bar in the Laurentian Great Lakes have been reported e.g. by Rodgers ( 1965). A theoretical study of the thermal bar in a quasi-steady state in a dimictic lake was presented by Huang (1972), whereas Brooks and Lick (1972) presented a time- dependent theoretical study of the development of the thermal bar. A review of the relatively few studies of cooling processes in temperate lakes combined with the observation of the evolution of thermal bars in two deep lakes is given by Carmack and Farmer (l 982). Recent field investigations of the thermal bar in Lake Ladoga using satellite data were conducted by Malm and Jonsson (1993) and Malm et al. (1994). A theoretical model of the thermal bar migration in a circular lake was presented by Zilitinkevich and Malm (1993). A description of the spring thermal bar in Lake Baikal was presented by Shimaraev et al. (1993). The authors suggest that the thermal bar at the south-eastern shore of the Central Basin initiates a thermobaric instability leading to a jet of cold water which apparently sinks from the hypolimnion on the open-lake side of the thermal bar to the deep part of the Central Basin. However, it remains unclear how the cold surface water is able to penetrate through the MTM. A detailed analysis of the thermal bar by Peeters et al. ( 1996) revealed that it can only explain vertical mixing in the top 200 to 300 m. During our field investigations on Lake Baikal between 1992 and 1995 we observed two different types of thermal bars. ( l) In spring, a river induced thermohaline front becomes established at the east coast of the Central Basin, separating the warm, more turbid and saline Selenga water near the shore from the clear and cold mid-lake water. (2) In autumn, a "pure" thermal bar which is determined exclusively by the evolution of the temperature stratification becomes established in the northern part of the Northern Basin and migrates slowly from the near-shore region to the central part of the basin. In the following, both types of thermal bar will be described in detail. 4.2 Spring Thermal Bar Fig. 4.1 shows the two-dimensional (6,Sc)-distribution obtained from CTD profiles. They where measured along a transect across the Central Basin between Boldakovo and Olkhon on June 8, 1993. Generally, the small horizontal temperature and ionic salinity gradients illustrate the remarkable horizontal homogeneity of the water between the eastern and the 55 Boldakovo 0 3.6- - 200 ---- -3.6------ 400 600 800 1000 1200 0 200 _-4 _____95.2 - .,.. .,.. 400 600 800 1000 1200 5 10 15 20 25 30 35 40 45 50 Distance from Boldakovo [km] Figure 4.1: Two-dimensional isopleths of (a) potential temperature 6 and (b) ionic salinity Sc measured along a transect between Boldakovo and Olkhon on June 12, 1993. See also Fig. 3.10. western shores. Exceptions can be seen in the slight increase in surface temperatures across the lake from the west (3.4 °C near Olkhon) to the east (3.7 °C near Boldakovo), and the distinct increase in ionic salinity near Boldakovo. The low surface temperatures near Olkhon are the result of the southerly flow of cold water from the Northern Basin along the western 56 shore (Shimaraev et al. 1994). The warm and saline water at the eastern shore originates from the Selenga. Note that within 2 km distance from the shore, temperatures already exceed 4 °C, and a thermal bar develops. Unfortunately, the coastal region is not included in the transect. The warm-water lens at a depth of 1000 m approximately 12 km offshore was already observed 2 weeks earlier on May 23 (Fig. 3.10). It consists of water from the Selenga, which, during early spring, flows along Kukui Canyon to the deep part of the Central Basin (see also Sect. 3.4.1 ). This section focuses on the temporal development of the thermal bar at Boldakovo in late May and early June. The thermal bar which develops at the east coast of the Central Basin in late spring is to a large extent controlled by the northerly flow of warm and saline Selenga water along the shore (Shimaraev et al., 1993). According to chemical analysis of water samples taken at Boldakovo from either side of the thermal bar and from the Selenga (L. Sigg, unpublished data), the water on the shore-side of the thermal bar is composed of 20% to 50% river water (Tab. 3.3). Since the river's salinity exceeds the lake's, a distinct salinity gradient is superimposed on the temperature gradient across the thermal bar. The stratification encountered at the east coast of the basin in spring is therefore not a pure thermal bar but rather a thermohaline front. For the sake of simplicity, however, in the following the term "thermal bar" will be retained. Central Basin Figure 4.2: Map of the east coast of the Central Basin at Boldakovo. The symbols ~ mark the positions where CTD profiles were .a measured: (•) at positions A, B and C on ~ 52°35' May 28 (Fig. 4.3). June 7 (Fig. 4.6) and i-1 June 16, 1993 (Fig. 4.8); (o) on May 30, (x) June 5 and (D) June 11, 1993 (Figs. 5 lOkln 4.10 and 4.11). The dotted lines are depth contours. 107°10' 107°20' 107°30' Longitude [0 E] During the spring 1993 expedition, numerous CTD profiles were measured near the shore at sampling locations shown in Fig. 4.2. CTD profiles taken 1 km offshore (pos. A in Fig. 4.2), 4 km offshore (pos. B), and 9 km offshore (pos. C) between May 28 and June 16, 1993, are presented in Figs. 4.3 to 4.9. 57 The profiles taken on May 28, 1993, are plotted in Fig. 4.3. A blow-up of the region between depths of 50 m and 250 mis given in Fig. 4.4. According to the temperature profile at pos. A (Fig. 4.3a), the thermal bar is located about 1 km offshore. The noisy signal in the uppermost 50 mis caused by water with different (8,Sc)-characteristics interleaving and illustrates the turbulent situation pertaining near the thermal bar. In the offshore region (pos. B and C), winter stratification still persists. In the uppermost 80 m, the water column is inversely stratified with respect to temperature. At pos. B, the MTM is located at a depth of about 100 m and does not coincide with Tmd· which is not reached until about 230 m depth (Fig. 4.4a). Because the temperature is close to Tmd. where a"" 0, the destabilising effect of the slight temperature decrease between depths of 100 m and 230 mis small. At pos. C, the temperature is extremely constant (3.59 °C) between a depth of 150 m and the point of intersection with the Tmd·line, which is located at a depth of about 190 m. The bottom temperature (at 1200 m depth) is 3.22 °C. Ionic salinity (Fig. 4.3b) measured at pos. Band C is large and quite homogeneous in the uppermost 50 m. In fact, it is almost identical with that measured at pos. A. Sc decreases markedly at depths between 50 m and 70 m and increases slightly thereafter. Fig. 4.4b illustrates that at pos. B, ionic salinity increases by about 0.1 mg kg-I between the MTM and the Tnut"line. At pos. C, Sc exceeds 96 mg kg-I between depths of 600 m and 900 m. The noisy signal below 500 m suggests that these relatively high values are the result of Selenga water flowing along the eastern shore into the deep part of the basin. 200 ] 400 i 600 0 800 b) 1000 1200 u.i-u.J.J'-'--U...... ,,-U..L-.u..J...... U.U 2.5 3 3.5 4 4.5 95 96 97 98 0.055 0.06 0.065 1 3 [mgkg" ) p • 1000 (kg m· ] 0 l°C1 s, qua Figure 4.3: (a) Potential temperature 8, (b) ionic salinity Sc and (c) quasi-density Pqua of CTD profiles measured off Boldakovo on May 28, 1993. The profiles were measured I km off shore (pos. A in Fig. 4.2: dashed line), 4 km off shore (pos. B: dotted line), and 9 km off shore (pos. C: solid line). The straight thin, dotted line shown in (a) is the temperature of maximum density, Tmd· 58 50 ...... !"":.., ·- :~ May 28, 1993 100 ....., ] \ .... <"' -:5 150 ~ .J ~ 0 r 1 ~ 200 a "J:.:;:r#·. 't., b) c ) .. ~ \ 250 ' 3.5 3.55 3.6 95.1 95.2 95.3 95.4 0.054 0.055 0.056 1 3 ) p -1000 (kg m" e [°C1 s. [mg kg" qua J Figure 4.4: Blow-up of the profiles taken at positions B and C between 50 and 250 m depth. Quasi-density Pqua (Eq. 2.22) calculated from potential temperature, ionic salinity, and from the silicic acid concentrations approximated by the mean background field (Tab. 3.2) on May 28 is shown in Fig. 4.3c. According to the distribution of Pqua• the stratification near the thermal bar at pos. A is neutrally stable or slightly unstable, which reflects the turbulent situation revealed by the noisy signals of e and Sc . At pos. B and C, quasi-density increases with depth, indicating that the water column is stably stratified. Maximum stability is reached between depths of 50 m and 70 m, where the increase in Pqua with depth is greatest. The stabilising effect of the temperature increase in the tbermocline is obviously compensating for the destabilising effect of the distinct salinity decrease. Between depths of 80 m and 250 m, Pqua is extremely constant, indicating that the various water masses with different ( 9,Sc)· characteristics are located at positions where they are neutrally stable. The ( 6,Sc)-diagram of the three profiles (Fig. 4.5a) reveals two different types of water. The profile at pos. A represents water with a relatively high concentration of dissolved ions (96.5 - 97 mg kg-1). Water with similar ionic salinity but lower temperature is found in the uppermost 50 mat pos. B and C. The deep water of the undisturbed water column in the mid-lake region (95.1 -95.3 mg kg-I) is represented by the bottom water (below 1150 m) at pos. C. Similar water is found in the region of the MTM at pos. B and C (70 m - 270 m, and 70 - 390 m, respectively). The remaining water masses can be explained as a mixture of these two end-members and by heating or cooling. An enlargement of the (6,Sc)-diagram showing the data between depths of 50 m and 250 m (Fig. 4.5b) illustrates that the noisy 59 I 4.5 • Pos.A - 3.6 6 Pos. B x Pos.C 4 ~ O' ,' O' ~ 3.5 • 3.55 ~ (I) (I) 3 x b) - 3.5 2.5 x I 95.0 96.0 97.0 95.1 95.3 1 s, [mg kg' ] Figure 4.5: (a) (8,Sc)-diagram of the profiles taken at positions A, B and C on May 28, 1993 (see also Fig. 4.3). (b) Blow-up of the profiles taken at positions Band C between depths of 50 and 250 m. temperature and ionic salinity signals at pos. B (Fig. 4.3) are the result of two water masses mixing. In contrast, the water at pos. C is remarkably homogeneous. The profiles taken on June 7 are plotted in Fig. 4.6. According to the temperature profiles (Fig. 4.6a), the thermal bar is located somewhere between pos. B and C. On the inshore side of the thermal bar, summer stratification develops. Surface temperature at pos. A exceeds 11 "C already, and a distinct temperature gradient has developed. In the offshore region of the thermal bar, the surface water is warmer than on May 28, but winter stratification endures. The uppermost 50 to 80 m are convectively mixed. At pos. B, the MTM is located at about 150 m depth (TMrM = 3.60 °C), i.e. about 80 m above the point of intersection of the temperature profile with the TmJ·Iine. At pos. C, the temperature is extremely constant (9= 3.59 °C) between 150 m depth and the point of intersection with the Tmd-line at about 190 m depth. The bottom temperature (at 1200 m depth) is 3.22 °C. The ionic salinity (Fig. 4.6b) measured at pos. A reveals a significant increase in Sc on the near-shore side of the thermal bar. The surface value exceeds 105 mg kg-I. Due to convection in the mid-lake region, Sc is homogeneous in the uppermost 50 to 80 m but slightly lower than on May 28. At pos. B, an intrusion of low-salinity water can be identified between 560 and 600 m depth. The salinity maximum detected at pos. C between 600 and 900 m depth on May 28 (Fig. 4.3b) no longer exists. 60 0.055 0.06 0.065 3 ] pqua • 1000 [kg m" Figure 4.6: As Fig. 4.3 with data from June 7, 1993. The development of a stable summer stratification at pos. A is reflected in the pronounced increase in Pqua with depth (Fig. 4.6c). The weakening of the thermocline in the mid-lake region is illustrated by a decrease in the gradient of Pqua between 50 and 80 m dep that pos. B and C compared to the situation encountered on May 28. Due to convection, Pqua is homogeneous in the uppermost 50 to 80 m. The ( 8,Sc)·diagram of the profiles measured on June 7 (Fig. 4. 7a) reveals two different water masses. Warm and saline Selenga water can be identified at the lake surface at pos. A The volume fraction of Selenga water (Sc,r= 126 mg kg-1: Votintsev, 1993) at pos. A (Sc= 106 mg kg-1), calculated from the signals in the ionic salinity using Eq. 3.14, is 11mix = 0.34 (where Sc,/= 95.5 mg kg-I is the mean ionic salinity in the Central Basin). This value agrees 12 • Pos.A •• 10 15, Pos.B • Pos.c ue... 8 a.> • 6 ;'• • 4 / Figure 4.7: As Fig. 4.5a with data from June 7, 1993. 95 100 105 1 s< [mgkg" ] 61 well with the volume fraction determined from chemical data from water samples taken at Boldakovo and from the Selenga (20% to 50%; Tab. 3.3). The deep water from the undisturbed water column is characterised by its low ionic salinity (Sc < 95. 3 mg kg-1). The ( O.Sc)-characteristics of all the remaining water is located along the mixing line between these two end-members. On June 16, the thermal bar is located approximately 9 km offshore (Fig. 4.8a). The temperature in the uppermost 50 m at pos. C is slightly lower than TmJ, whereas at pos. B, the surface temperature exceeds Tmd· Below a depth of 50 m ( 8 > Tmd ), a weak stratification develops. An intrusion of warm water can be identified at pos. C at about 300 m depth. In the near-shore region at pos. A, the surface temperature exceeds 14 °C. The ionic salinity in the surface water at pos. A exceeds 110 mg kg-1 (Fig. 4.8b). Further off shore, at pos. B and C, Sc is significantly higher than on June 7. The Sc-signal at pos. C is disturbed by numerous intrusions. At about 300 m depth, a pronounced signal of high salinity water can be identified which coincides with the increase in 8 apparent in Fig. 4.7a. Below 300 m depth, the ionic salinity signal is highly variable most probably as a result of the flow of Selenga water into the deep part of the basin. The large increase in quasi-density at pos. A illustrates that in the near-shore region, the summer stratification has intensified (Fig. 4.8c). In the mid-lake region at pos. B and C, where temperature is close to Tmd· quasi-density is homogeneous in the uppermost 80 to 100 m, indicating that the stratification is neutrally stable. Despite the intrusion of warm, 0.055 0.06 0.065 3 p - 1000 [kg m· ) qua Figure 4.8: As Fig. 4.3 with data from June 16, 1993. 62 high salinity water at pos. C at a depth of about 300 m, the slight increase in Pqua below 100 m depth illustrates that the stratification is weak but stable. Note that according to the gradient of Pqua the intruding water would be subject to positive buoyancy. 14 • 12 • O' 10 ~ a:> 8 • , • Pos.A Pos. B 6 A Figure 4.9: As Fig. 4.5a with data from June x Pos.C 16, 1993. 4 ' 95 100 105 110 1 S< [mgkg' ) The ( e.Sc)-distribution of the three profiles taken on June 16 (Fig. 4.9) is very similar to the one shown in Fig. 4. 7. The volume fraction of Selenga water at the lake surface at pos. A (Sc"" 112 mg kg-I) calculated from ionic salinity (Eq. 3.14) is T/mix = 0.54. The water from the undisturbed water column (Sc< 95.3 mg kg-I) is represented by the deep water below a depth of 1100 mat pos. C. At pos. B, the ionic salinity has increased slightly and ranges from 95.4 to 96.6 mg kg-I. To illustrate the temporal development of the ( 8,Sc)-distribution near the thermal bar, two-dimensional isopleths obtained from profiles taken along short transects at Boldakovo on May 30, June 5 and June l l, 1993, are shown in Figs. 4.10 and 4.11. Note that sampling dates and positions are different from those of the profiles shown in Figs. 4.3 to 4.9. The isotherms (Fig. 4.10) illustrate the faster increase in surface temperature near the shore than in the mid-lake region. Note that the profiles near the shore were not taken at exactly the same positions (Fig. 4.1), e.g. the transect measured on June 11 begins closer to the shore than the transects measured on May 30 and June 5. Due to the temperature increase in the surface layer, the vertical temperature gradient intensifies, which causes the stability of the summer stratification to increase. In the mid-lake region, winter stratification gradually weakens during the observation period as the water temperature approaches Tmd· The horizontal temperature gradient between the warm water in the coastal region and the cold water in the central part of the basin results in vertical isotherms in the region between 1 km and 3 km offshore. The position where water temperature is equal to the local Tmd is plotted 63 as thick dotted line in Fig. 4.10. The T md-line separates water with () > Tmd from water with fJ < T md. At the lake surface the Tmd-line represents the position of the thermal bar; below the lake surface, it represents the MTM. Due to the poor horizontal resolution of the profiles taken on May 30 (Fig. 4. lOa), the position of the Tmd'"line cannot be reliably determined and therefore it is not plotted at the lake surface. From Fig. 4.3 we can conclude though that the thermal bar is located approximately 1 km offshore. The Tmd-line descends from a depth of about 50 m l km offshore to the MTM (TMTM"" 3.58 °C) located at about 220 min the offshore region. As indicated by the 3.6 °C isotherm, relatively warm water migrates along the Tmd-line towards the central part of the basin, causing the temperature of the MTM to increase slowly. On June 5 (Fig. 4. lOb), the thermal bar is located approximately 1.5 km offshore. The Tmd-line is nearly vertical in the top 150 m and then leans horizontally towards the central part of the basin. The MTM (TMTM"" 3.60 °C) is located at about 200 m depth. On June 11 (Fig. 4. lOc) the thermal bar is located about 2 km offshore. The Tmd-line runs between the 3.8 °C and the 3.9 °C isotherms and does not sink deeper than 70 m depth. Due to the increasing temperatures in the central part of the lake, the MTM has almost disappeared (TMTM > 3.9 °C). The development of the Sc-distribution between May 28 and June 11 is shown in Fig. 4.11. The isohalines illustrate the distinct salinity gradient between the saline Selenga water near the shore and the less saline mid-lake water. Note that despite its large concentration of dissolved ions, the near-shore Selenga water is too warm to sink to the bottom of the Central Basin. At 8 = 4. 5 °C, a temperature difference of about 0.5 K is sufficient to compensate for the density difference caused by the salinity difference of 5 mg kg-I (/3c = 0.8107·10-3 (g kg-1)-1). According to Fig. 4.10, the lateral temperature difference between the shore and mid-lake is always larger than 2 K. At the lake surface, the isohalines reveal a layer of water with relatively large salinity (96 mg kg-I) suggesting that on the offshore side of the T md-line, small portions of Selenga water are also present. In the offshore region, Sc lies between 95.5 and 96.5 mg kg-I. Due to the flow of saline inshore water towards the central part of the basin, Sc increases steadily between 3 km and 4 km offshore, as illustrated by the 96 mg kg-l isohaline. On June 5, water with Sc =96.0 mg kg-I is intruding into the open water column at about 500 m depth (Fig. 4.8b). On June 11, the ionic salinity 3.8 km offshore below 100 m lies between 96.0 and 96.5 mg kg-1. Neutral tracks (Peeters et al., 1996) calculated for specific water parcels, and neutral surfaces {McDougall, 1984, 1987a) for the situation encountered on June 5, 1993, are plotted in Fig. 4.12. The neutral tracks {thick solid lines) give the direction of neutrally buoyant transport of a water parcel that does not mix with the ambient water over a finite distance. In contrast, the neutral surfaces (short lines) represent the direction of neutrally buoyant displacement of a parcel over an infinitesimal distance (See Sect. 2). Both neutral tracks and 64 0 100 200 ! -:5 300 a.. QI 400 0 500 600 a) May 30, 1993 0 100 200 ! -:5 300 a.. QI 400 0 500 600 b) June 5, 1993 0 iii 117•• i] •• 100 ---=::=3.9 200 ~-----3.8- ! 300 o:f3 a.. QI 0 400 500 600 c) June 11, 1993 700 0 1 2 3 4 Distance from shore [km] Figure 4.10: Two-dimensional isopleths of potential temperature 9 along a short transect offshore of Boldakovo measured on (a) May 30, (b) June 5, and (c) June Jl, 1993. See Fig. 4.2 for positions of stations. The position of the local Tmd is indicated by the thick dotted line. 65 0 100 200 ] 300 ii Q, Q; 400 Q 500 600 0 100 200 ] 300 iiQ, Q; 400 Q 500 600 0 100 200 ] ii 300 ~ Q 400 500 600 700 0 1 2 3 4 Distance from shore [km] Figure 4.11: Two-dimensional isopleths of ionic salinity Sc along a short transect offshore of Boldakovo measured on (a) May 30, (b) June 5, and (c) June 11, 1993. See Fig. 4.2 for positions of stations. 66 neutral surfaces are the result of a delicate interplay between salinity and temperature, and allow some characteristic features of the spring thermal bar encountered at Boldakovo to be described and interpreted. According to the neutral tracks, neither the warm and saline Selenga water inshore (e) nor the cold mid-lake surface water (©) has the potential to sink to the deep part of the basin. The former is too warm ( (J > Tmd) and, despite its high salinity, not dense enough to overcome the stable summer stratification that develops near the shore. The latter is too cold to penetrate through the MTM, and remains on the offshore side of the thermal bar. To maintain the cold bottom temperature found in the Central Basin (Fig. 3.4c), cold surface water ( (J < 3.1 °C) must mix downwards. Surface temperatures this low are found on the offshore side of the thermal bar only. Since salinity gradients are extremely small in this region, the neutral track (©) is predominately determined by the temperature stratification (i.e. it follows the isotherms) and does not sink deeper than 80 m depth. Intense vertical exchange is only observed close to the thermal bar. According to its neutral track, surface water with temperatures close to Tmd (0) may sink to a depth of about 300 m. In contrast to the neutral tracks, the neutral surfaces (short lines in Fig. 4.12) represent the direction of flow of water parcels that are continually mixed with the ambient water. At Boldakovo both neutral tracks and neutral surfaces are mostly parallel and define the narrow range within which the water is expected to move. 0 100 200 ! ofj 300 Q., QI Q 400 500 600 0 1 2 3 4 Distance from shore [km] Figure 4.12: Neutral tracks (thick solid lines) and neutral surfaces (short lines) at Boldakovo for the situation encountered an June 5, 1993. The symbols mark the initial position of different neutral tracks. Some isotherms and isohalines from Figs. 4. lOb and 4.11 b are included. 67 The observations presented in Figs. 4.3 to 4.12 can be combined to the following scenario. The spring thermal bar at Boldakovo appears in mid May and lasts until mid June. It is to a large extent influenced by the flow of warm, and saline Selenga water along the eastern shore at the lake surface. In fact, according to the ionic salinity, the near-shore surface water is composed of 30% to 55% river water. But also in the surface layer in the mid-lake region small amounts of Selenga water can cause a slight increase in ionic salinity. During the early stage of the thermal bar, the warm and saline Selenga water is trapped near the shore by the cold and less saline lake water. The northerly flow of the river water implies that the transport of heat is mainly parallel to the shore. Consequently, the thermal bar migrates slowly off shore, as spring heating continues and the discharge of the Selenga increases. Between the end of May and the middle of June it moves from about 1 km to 9 km offshore. Vertical exchange in the thermal bar region is restricted to the uppermost 300 m. 4.3 Autumn Thermal Bar A different type of thermal bar was encountered in the Northern Basin in autumn 1994. Compared to the thermohaline front at Boldakovo, the autumn thermal bar develops in the opposite direction. Water temperature near the shore drops below 4 °C while the mid-lake surface water is still warmer than 4 °C. As autumn cooling progresses, the temperature of the surface water decreases and the cold water front moves offshore towards the central part of the basin. The thermal bar was first observed in the shallow northern part of the Northern Basin on October 28, 1994 (Fig. 4.13). Two-dimensional isopleths of potential temperature and ionic salinity measured along a short transect across the thermal bar are shown in Fig. 4.14. The profiles were taken within a two-hour period (from 11 am to 1 pm) at the positions shown in Fig. 4.13. The isotherms (Fig. 4.14a) reveal a rather dynamic situation. The Tmd- line (thick dotted line) separates the cold inshore water ( 8 < Tnut) from the wanner offshore water ( (} > T md ). A second branch of the T md-line is located in the offshore region below lOOmdepth. In contrast to the thermohaline front at Boldakovo, the thermal bar observed in the Northern Basin is a "pure" thermal bar, i.e. it is predominately determined by the evolution of the temperature stratification. The isohalines in Fig. 4.14b indicate that horizontal salinity gradients hardly exist and vertical salinity gradients are small. 68 55°45' 45 /?....z .,, 109°36'15" 109°36'45" Longitude [0 E] Figure 4.13: (a) Map of the northern part of the Northern Basin. The symbols mark the positions where CTD profiles were taken (•) on October 28 and (o) on November 8, 1994. The Upper Angara (A) is the major river entering the basin. A blow-up of the positions of the measurements on October 28 is shown in (b). The shaded area marks the near-shore region where surface temperatures are below 4 °C. Fig. 4.15 shows a ( 8,Sc)-diagram of some of the profiles included in Fig. 4.14. The profiles measured in the offshore region at positions 49 and 50 have similar (8,Sc)-properties and represent water from the undisturbed water column in the northern part of the Northern Basin. Ionic salinity ranges from 94.l to 94.7 mg kg-1 and temperature from 3.55 to 4.1 °C. The profiles measured at positions 43 and 45 reveal that (1) the bottom water becomes wanner and slightly less saline and (2) the salinity range decreases on moving towards the shore. The increases in bottom temperature and decrease in ionic salinity suggest that surface water with 8"" Tmd and with a slightly lower concentration of dissolved ions was convectively mixed with the water from the undisturbed water column below 100 m depth. Convection near the thermal bar results in a rather homogeneous water mass. Such water is represented by the profile taken at position 42. The data give no evidence that the creation of the thermal bar in the coastal region is initiated by the westerly flow of cold and less saline water from the Upper Angara (Sc= 81.3 mg kg-I: Votintsev, 1993) along the northern shore of the basin. Although the inflow from the Upper Angara during October and November is still relatively large (100 - 200 m3 s-1: Shimaraev et al. 1994), no water with significantly different (8,Sc)-characteristics can be identified in Fig. 4.15. 69 North South 0 40 ] -;$ 80 Q"'"" 120 a) (J [°CJ 160 47 49 50 0 40 94.2 ] t 80 Q"" 120 '~~·4'_)4. 1 b)Sclmgkg' ) 94.6 160 0 0.2 0.4 0.6 0.8 1.0 1.2 Relative Distance [km] Figure 4.14: Two-dimensional isopleths of (a) potential temperature (Jand (b) ionic salinity Sc along a short transect in the near-shore region of the northern part of the Northern Basin on October 28. 1994. The position of the local Tmd is indicated by the thick dotted line in (a). The numbers between the figures mark the station numbers where the profiles were taken (See Fig. 4.13). 70 4.1 42 0 43 4 0 45 x 49 3.9 50 ua.. Cl) 3.8 0 + 3.7 8 3.6 Figure 4.15: (0,Sc)-diagram of profiles taken • along a transect across the thermal bar on 94.1 94.3 94.5 94.7 October 28, 1994. See also Fig. 4.14. 1 Sc [mg kg- ]'+ Two weeks later, on November 8, the thermal bar was observed 40 km further south (Fig. 4.13) in the central part of the basin. The two-dimensional (8,Sc)-distribution obtained from profiles taken along a transect across the thermal bar are shown in Fig. 4.16. The Tm,;- line (thick dotted line) extends horizontally over a distance of about 5 km, and vertically from the water surface to about 150 m depth. It separates the cold-water region ( 8 < T md ), where the winter stratification develops, from the warmer mid-lake region ( 8 > Tmd ), which is convectively mixed over the upper 180 m. The temperature of the convective layer is approximately 4.1 "C. Only water temperatures in the uppermost 30 mare colder implying that the surface layer is slightly unstable. Note however that due to the surface temperatures being close to T md· where a= 0, the destabilising effect of the inverse temperature stratification is very small. Supported by wind shear and/or further cooling, the temperature inversion at the lake surface will lead to convection resulting in the erosion of the 4.1 °C isotherm, as can be observed near the Tmd·line. Subsequent to the erosion of the 4.1 °C isotherm, the thermal bar will migrate towards the central part of the basin. The salinity distribution is shown in Fig. 4.16b. In the 180 m convective surface layer, salinity is low and homogeneous (93.7 mg kg--1), but increases markedly between 180 and 200 m depth. On the cold-water side of the Tmd·line, this distinct salinity increase is not observed. Here, the concentration of dissolved ions is intermediate and rather homogeneous (94.3 94.4 mg kg-I) in the top 200 m. The small gradient in ionic salinity across the thermal bar results in vertical isohalines. 71 North South 0 100 ] 200 o£j 0.. ~ Q 300 400 500 86 87 88 92 91 90 , 81 0 ;- I I \ , - ---_, (- - ' I 93.1 100 94.3 ] 200 94.4 ,- ·94.5 - - : ' t~ Q 300 1 400 b) Sc [mg kg- 1 """"------"'.:. .... - ::-;.:_ -..:- -94.7' ·, 500 0 1 2 3 4 5 6 7 8 9 10 Relative distance [km] Figure 4.16: Two-dimensional isopleths of (a) potential temperature 6 and (b) ionic salinity Sc along a transect in the northern part of the Northern Basin on November 8, 1994. The numbers between the figures mark the station numbers where the profiles were taken (See Fig. 4.13). 72 • 86 4.1 " 87 0 88 4 x 91 -u 3.9 od" ...___+__ ,_o J' I:- 3.8 ~Q) 0Q x e <:f 3.7 3.6 Figure 4.17:(9,Sc)-diagram of profiles taken !~\ along a transect across the thermal bar on 3.5 : November 8, 1994. See also Fig. 4.16. 93.8 94.2 94.6 1 Sc [mgkg- ] A (6,Sc)-diagram (Fig. 4.17) of some of the profiles included in Fig. 4.16 illustrates the changes in temperature and ionic salinity occurring across the thermal bar in the uppermost 250 m. Due to convective mixing near the thermal bar, the concentration of dissolved ions becomes homogeneous, i.e. the range in ionic salinity decreases. In the profiles measured on the offshore side of the Tmd-Iine at positions 90 and 91, Sc ranges from 93.7 to 94.7 mg kg-I. In the profiles taken on the near-shore side of the TmJ-Iine at positions 86 and 87, Sc lies between 94.3 and 94.7 mg kg-I. Fig. 4.17 also confirms that the convective mixing near the thermal bar is restricted to the region from the surface to 250 m depth. Below 250 m, where Sc> 94.5 mg kg-I, the profiles reveal no significant differences in temperature or ionic salinity across the Tmd"'line. To visualise the dynamic structure near the thermal bar, neutral surfaces and neutral tracks are plotted in Fig. 4.18. On the near-shore side of the Tmd-line, a weak winter stratification is developing. According to its neutral track, cold surface water with 6 > Tm&(e) is not able to sink below the Tmd-Iine. On the open lake side of the thermal bar, the surface water is slightly colder than the isothermal layer bounded by the 4.1 °C isotherm, and the water column is slightly unstable. The neutral track of a surface water parcel (0) leads along the 4.05 °C isotherm and is unstable in the uppermost 50 m, i.e. it separates denser water above from less dense water below. The parcel will therefore not move along its track, but will sink by free convection. At about 200 m depth, it will be intercepted by its neutral track, which at this depth is stable. An additional neutral track illustrates that in the region where 6 =T md (0), the water column is vertically unstable. The horizontal mixing of water masses from either side of the thermal bar results in a density increase and may lead to local vertical instabilities in the water column (cabbeling). Since the concept of neutral tracks does not include the effect of mixing between the water parcel and with the ambient water, it cannot adequately account for such mixing processes. However, the resulting vertical instabilities can be identified by the neutral track. The water parcel with 6 = T md (0) is 73 North South 0 100 ] ..s 200 e- 0 300 400 3. - 500 --.... -- 0 1 2 3 4 5 6 7 8 9 10 Relative distance [km] Figure 4.18: Neutral tracks (thick solid lines) and neutral surfaces (small lines) in the northern part of the Northern Basin on November 8, 1994. The symbols mark the initial position of different neutral tracks. Some isotherms and isohalines from Fig. 4.17 are included. located above its track, i.e. it represents water that is denser than the ambient water. The parcel will consequently sink to its neutral track, which is located about 150 m below. In Fig. 4.18, the neutral surfaces (short lines) and neutral tracks are nearly parallel to each other and define the narrow depth range within which the water masses are expected to move. While the neutral tracks do not penetrate deeper than 200 m, the neutral surfaces point more or less in a horizontal direction below 250 m depth. Both complementary concepts confirm that only the uppermost 250 m of the water column are influenced by the convective mixing related to the thermal bar. Although it is not possible to present a detailed description of the development of the autumn thermal bar in the Northern Basin, the observations presented in Figs. 4.13 to 4.18 can be combined to the following scenario. In late October, the autumn thermal bar becomes established in the near-shore region of the shallow northern part of the Northern Basin. From its initial position it moves rapidly offshore. Two weeks later, it is already located about 40 km further south in the central part of the basin. As the thermal bar migrates offshore, the surface layer near the Tmd-line is mixed convective!y. However, the vertical exchange is restricted to the top 250 m. Below, temperature and ionic salinity do not change. Although the inflow from the Upper Angara is relatively large in October and November, the development of the thermal bar in the Northern Basin is not influenced by the river. 74 4.4 Summary and Conclusion In Lake Baikal, two different types of thermal bars occur. In spring, a river induced thermo- haline front establishes at the east coast of the Central basin. It separates the warm and saline water near the shore from the clear and cold mid-lake water. The inshore water consists between 30% and 55% of water originating from the Selenga. The northerly flow of river water provides a continuous supply of warm water, implying that heat is mainly transported parallel to the shore and just to a small extent towards the central part of the basin. Within the two week observation period, the thermohaline front does not significantly move off shore towards the central part of the basin. In autumn, a "pure" thermal bar that is entirely determined by the evolution of the temperature stratification establishes in the northern part of the Northern Basin. Horizontal salinity differences are small and are the result of convective mixing processes near the thermal bar. In contrast to the thermohaline front at Boldakovo, the thermal bar migrates fast towards the central part of the basin. Within 2 weeks, it moved about 40 km off shore. Generally, the importance of the thermal bar phenomenon for the deep-water ventilation in Lake Baikal is overestimated. Intensive vertical exchange occurs only in the region near the thermal bar where the water column is unstable. However, the sinking surface water is too warm to plunge to the deep part of the lake. In the Northern Basin, only the water column in the top 250 mis turned over by convective mixing related to the thermal bar. Supported by the enhanced salinity, the water near the thermohaline front at Boldakovo sinks to about 300 m depth. Disturbances in the ionic salinity signal in the Central Basin below 300 m are most probably caused by the saline water from the Selenga flowing in northerly direction in the deep part of the basin along the eastern shore. 5. Tritium and Noble Gas Analysis: Concepts and Data Processing Starting in the early fifties, large amounts of tritium (3H) were injected into the atmosphere by thermonuclear weapon tests conducted by the USA, USSR, UK and later on by France and the People's Republic of China. Within a few years, the global tritium inventory increased by more than a factor of 100 (Weiss and Roether, 1980). This most doubtful episode of human history unwittingly provided a powerful tool for studies of water movement which has since found numerous applications in oceanography, hydrology and limnology. 3H is the radioactive isotope of hydrogen and decays to the stable helium isotope, 3He. Its half-life of 12.43 years provides an ideal time-scale to assess large-scale mixing processes in lakes. The 3H-3He dating method was introduced to oceanography by Jenkins and Clarke (1976) and to limnology by Torgersen et al. (1977). Compared with its use in oceanography, relatively little effort has been made to use the method for lake research on a routine basis. A brief review of applications of the 3H-3He method in lacustrine systems is given in Sect. 6.1. In the following, a brief introduction to the 3H-3He dating method will be given. For technical details the reader is referred to Kipfer (1991). Details on the calculation of 3H-3He water ages are described by Aeschbach-Hertig (1994). A review of the method's applications in oceanography, hydrology and limnology is given e.g. by Schlosser (1992). 5.1 Geochemical Background 5.1.1 Tritium Tritium, the radioactive hydrogen isotope of mass 3, decays to the stable helium isotope 3He by ~-decay. Its half-life is 12.43 yr (Unterweger et al., 1980). Tritium concentrations are 76 generally given in tritium units (TU): 1 TU is equivalent to l tritium atom per 1018 hydrogen atoms. Tritium is produced naturally, mainly in the upper atmosphere, by the interaction of cosmic rays with nitrogen and oxygen: The mean natural production rate has been estimated to be 0.25 ± 0.08 3H cm-2 s-1 (Craig and Lal, 1961; Rozanski et al., 1991). About two thirds of the production takes place in the stratosphere and one third in the troposphere (Rozanski et al., 199 l ). After oxidation to HTO ("heavy water"), tritium enters the hydrological cycle by precipitation. The natural tritium concentration in continental precipitation is about 5 TU (Craig and Lal, 1961; Roether, 1967). The steady-state global inventory of natural tritium in the atmosphere is approximately 3.6 kg (Rozanski et al., 1991). The global tritium inventory was drastically increased by tritium produced in the atmosphere by the testing of thermonuclear weapons starting in the early fifties. Between 1952 and 1962, approximately 600 kg of tritium were injected into the atmosphere from tests carried out by the USA, USSR and UK (Rozanski et al., 1991). An additional 20 kg were added by tests conducted by the People's Republic of China and France between 1967 and 1980. After 1980, no further atmospheric thermonuclear tests were carried out. Most of the bomb-generated tritium was injected into the stratosphere, subsequently re- entering the troposphere primarily via tropospheric folding at mid and high latitudes in the northern hemisphere. Once in the troposphere, HTO is rapidly removed by precipitation and vapour exchange (Rozanski et al., 1991 ). Monthly tritium concentrations in precipitation from a network of about 250 long-term monitoring stations spread around the globe are available from the World Meteorological Organisation/International Atomic Energy Agency (WMO/IAEA; IAEA, 1992). The longest record comes from Ottawa (Canada) and started in 1953. The development of the tritium concentration in precipitation with time at four different locations is shown in Fig. 5.1: Ottawa, Canada (continental station in the northern hemisphere) Valentia, Ireland (marine station in the northern hemisphere) Manaus, Brasil (continental station in the equatorial region) Kaitoke, New Zealand (marine station in the southern hemisphere) 77 ---+- Ottawa 1000 -<>--Valentia Figure 5.1: Tritium concen- --··•····Manaus tration in precipitation at four ····o-·--Kaitoke representative stations around § 100 the globe: Ottawa (continental station, northern hemisphere); ;: Valentia (marine station, 10 northern hemisphere); Manaus (tropical station); Kaitoke (marine station, southern hemi· sphere). 1960 1970 1980 1990 Year The global distribution of tritium is highly asymmetric. Tritium concentrations in precipitation are 1 - 2 orders of magnitude lower in the southern hemisphere than in the northern hemisphere. The main reasons for this are as follows: (1) major bomb tests were carried out predominantly in the northern hemisphere at high latitudes (Rozanski et al., 1991); (2) interhemispheric exchange is impeded by the fact that the circulation in the stratosphere is directed upwards and polewards in the tropics, and downwards in the extratropics (Holton et al. 1995); (3) tropospheric moisture in the southern hemisphere originates predominantly from the ocean and therefore exhibits low tritium concentrations (Rozanski et al., 1991 ). Furthermore, tritium concentrations at continental stations are enhanced by a factor of 2 - 4 compared to marine stations at the same latitude, mainly because of precipitation rates over continents are lower than over the oceans and because removal of HTO by vapour exchange over the ocean is more effiecient (Doney et al., 1992). In the northern hemisphere, tritium concentration in precipitation reached its maximum in 1963, immediately after the first major atmospheric thermonuclear explosions (Fig. 5.1 ). Until 1967 it decreased exponentially at a mean rate of 0.57 ± 0.04 yrl (Rozanski et al., 1991 ). Subsequent to the French and Chinese tests, the tritium concentration remained almost constant between 1968 and 1971. Since 1971, the concentrations decreased at a rate of approximately 0.10 yrl in Valentia, Manaus and Kaitoke and approximately 0.07 yrl in Ottawa. The decline was disturbed by small inversions due to further testings after 1971. In the southern hemisphere, tritium concentrations and gradients are much smaller. The maximum value was reached in 1964, one year later than in the northern hemisphere. In recent years, tritium levels in precipitation have returned to values close to the natural pre-bomb values in most parts of the world (Rozanski et al., 1991). In some regions, the 78 concentrations are still significantly enhanced, for instance in Ottawa. Because of the generally low values, anthropogenic emissions of tritium, e.g. from nuclear power plants or from manufacturing industries, are becoming locally and regionally more and more visible. However, compared to the tremendous input from atmospheric nuclear weapon tests, these emissions are negligible. According to an estimate by Rozanski et al. (1991) the average global tritium discharge into the atmosphere by nuclear power plants in 1988 was about 24 g yrl. Recent concentrations in northern hemisphere lakes range from 10 to 30 TU. 5.1.2 Helium Helium has two stable isotopes: 3He and 4He. Helium is an extraordinary element in many ways. After hydrogen, it is the most abundant element in the universe. As a noble gas, it hardly participates at all in chemical reactions. Since helium is very mobile, it is an ideal tracer for geological migration processes. In terrestrial materials, the 3Hef4He ratio varies over five orders of magnitude, ranging from 5· 1Q--5 (in Hawaiian phenocrysts and nodules: Mamyrin and Tolstikhln, 1984) to 5· I0--10 (released from uranium minerals: Mamyrin and Tolstikhin, 1984). Due to the distinct signatures of helium in its main terrestrial reservoirs, the analysis of helium isotopes provides an instrument to identify the origin of gas fluxes from the Earth's interior (e.g. Clarke et al., 1969; O'Nions and Oxburgh, 1983; Torgersen and Clarke, 1985; Kipfer et al., 1994; Aeschbach-Hertig et al., 1996b). He in the Earth's Interior: Radiogenic 4He is produced in the Earth's crust by the radioactive decay of U and Th. The production rate in the crust is approximately 3· IQIO 4He atoms cm-2 s-1 (O'Nions and Oxburgh, 1983, Torgersen and Clarke, 1985). Nucleogenic 3He is produced indirectly by different nuclear processes via tritium, mainly by 6Li (n,a) 3H (fr) 3He (Morrison and Pine, 1955). The crustal helium isotope ratio is therefore highly dependent on the Li, U and Th content of the crust. Values of the crustal 3He/4He ratio (Rcr) range from 10--7 to 10--8. Radiogenic helium from the stable continental crust has a typical 3He/4He ratio of about Rcr"' 2·10--8 (Mamyrin and Tolstikhin, 1984). Mantle helium was first detected in the ocean (Clarke et al., 1969), where it is released mainly along mid-ocean ridges and has accumulated over time. It is interpreted as the remains of solar primordial He with a 3He/4He ratio of about I0--4 that was trapped during the formation of the earth (Mamyrin and Tolstikhln, 1984). Mantle helium seems to be the major reservoir of the light helium isotope in the earth. The best available estimate of the 3He flux 79 from the mantle into the sea, calculated from the measured helium excess in the ocean and the estimated deep-water renewal rate, is 3 ± 1 3He cm-2 s-1 (Craig et al., 1975, Welhan and Craig, 1983). Estimates of the mantle helium flux across the continental crust from accumulation rates in lakes are of the same order of magnitude (Sano et al., 1986). Mantle helium is characterised by a 3He/4He ratio Rm of about 3· l0-5 (Mamyrin and Tolstikhin, 1984). }lelium from the depleted upper mantle observed in Mid-Ocean Ridge Basalts (MORB) has a quite uniform helium isotope ratio ranging from 1.0· 10-5 to 1.2· I0-5 (Craig and Lupton, 1981 ). Helium from the undepleted lower mantle with a 3He/4He ratio exceeding the average MORB value by a factor of 2 - 3 is found in so-called "Hot Spots", e.g. Yellowstone or Kilauea (Mamyrin and Tolstikhin, 1984). A further mantle helium component with an isotope ratio of 0. 7 · l 0-5 to l.1·10-5 is found in subduction zones, e.g. in Kamchatka, in the Kuriles, and in New Zealand (Craig and Lupton, 1981 ). He in the Atmosphere: In the atmosphere, 3He is produced by the decay of tritium and, like tritium, by the interaction of cosmic rays with nitrogen and oxygen. The total net production rate was estimated by Craig and Lal (1961) as 0.5 3He atoms cm-2 s-1. The atmospheric 4He production is not significant. Both helium isotopes escape from the terrestrial atmosphere into space because of their low atomic masses. Thus, the atmospheric concentration represents an approximate balance between in situ production, input from the Earth's interior (from where He is continuously released), and escape from the upper atmosphere into space. The volumetric helium concentration in the atmosphere is 5.24 ± 0.05 ppm (Ozima and Podosek, 1983). The isotopic ratio of atmospheric helium is Ratm = 1.384· 10-6 (Clarke et al., 1976). Considering the long atmospheric residence time of 106 yr (Lupton, 1983), the assumption of a constant atmospheric isotopic ratio is reasonable over short periods of time. He in Natural Waters: Helium dissolved in natural waters originates from various different sources: The major part of the dissolved helium originates from the atmosphere. Equilibrium concentration, on the order of 5· I0-8 cm3STP g-1, is reached at lake surfaces by gas exchange. Since 3He is slightly less soluble than 4He, the helium isotope ratio in water is smaller than in the atmosphere (Req = 1.36· I0-6 for T = 4 °C: Benson and Krause, 1980). 80 An atmospheric helium excess (Ratm = 1.384· l Q-6) may be caused by complete dissolution of air (e.g. injected air bubbles). Helium from the mantle and/or the crust may be injected into the water through the sediment-water interface. From the helium isotope ratio it is possible to distinguish between helium of different geotectonic origins. 4He production by the a-decay of dissolved U and Th in the water column is negligible due to the low concentrations and to the extremely long half life of these isotopes. Excess 3He is produced by the decay of tritium. The tritiogenic helium component changes the 3He/4He ratio in favour of 3He. To summarise, the measured 3He and 4He concentrations in a water sample can be written as: (5. la) (5.lb) with : iHe101 measured helium concentration; iHe eq: equilibrium concentration in water; iHeex: atmospheric helium excess; iHe,.,: helium from the Earth's interior (mantle and/or crust); iHe1,;: helium produced by the decay of tritium. The different components can be distinguished by their typical 3He/4He ratios, summarised in Tab. 5.1. Table 5.1: Helium isotop ratio in different terrestrial compartments (R = 3Hel4He). 3Hef'IHe [10-6J RIRatm Reference Atmosphere 1.384 Ozima and Podosek (1983) Crustal production 0.02 0.015 Mamyrin and Tolstikhin (1984) MORB 11-14 8 -10 Craig and Lupton (1981) Hot Spots 20-40 15-30 Mamyrin and Tolstikhin (1984) Subduction Zones 7-11 5-8 Craig and Lu11ton ( 1981) 81 5.1.3 Neon Excess neon in water is indicative of atmospheric contamination. Neon is therefore analysed simultaneously with helium to identify the atmospheric helium component. Neon has three stable isotopes, 20Ne, 21 Ne and 22Ne, of which 20Ne is the most abundant (20NeJNero1 =0.9050 ± 0.0007; Ozima and Podosek, 1983). The atmosphere is the main terrestrial neon reservoir with a volumetric neon concentration of 18.18 ± 0.04 ppm (Ozima and Podosek, 1983). Neon is produced in the Earth's crust (e.g. Kennedy et al., 1990). There are indications of a solar primordial neon component in the Earth's interior (Honda et al., 1991; 1993). However, since the crustal neon production is much smaller than the 4He production 21 7 21 2 ( NdHe = 0.46± 0.08· 10- ; Nei2 Ne = 0.47 ±0.01; the crustal production of 20Ne is not significant; Kennedy et al., 1990), it is generally assumed that the atmosphere is the only source of dissolved Ne in lacustrine systems (Schlosser, 1989; Aeschbach-Hertig, 1994). Neon concentrations in a water sample therefore consist of an atmospheric solubility concentration and a possible atmospheric excess component: (5.2) 5.2 Calculation of the JU.JHe Age From the tritium and tritiogenic (i.e. produced by the tritium decay) helium concentrations, the 3H.3He age of a closed, homogeneous body of water can be calculated using the following basic equation (Torgersen et al., 1977): 3 T= l In ( l+ He3 ~" ·) , (5.3) with .t= 0.05576 yrl the decay constant of tritium (Unterweger et al., 1980), 3H the tritium concentration and 3He1,; the tritiogenic helium concentration (1 cm3STP g-l =4.019· 1014 TU). The 3H.3He age denotes the time since the water mass was last in contact with the atmosphere. For its calculation it is necessary to accurately determine the tritiogenic 3ffe component. From Eq. (5.1 b) it follows that: (5.4) 82 Introducing the specific helium isotope ratios of the different components (R = 3He/4He) and setting Rex= Ratrn yields (5.5) Since the atmosphere is assumed to be the only source of neon, the atmospheric excess 4Heex can be calculated from the neon excess (expressed as 20Ne ): 4 (20 20 ) 1 He,"= Ne101- Ne,q ·--, (5.6) Nairn where Narm = (20Ne/4He)airn = 3.14 (Ozima and Podosek, 1983). Next to the atmospheric helium excess, the terrigenic component accounts for the remaining 4He excess: (5.7) Inserting (5.6) and (5.7) into (5.4) yields the following expression (Aeschbach-Hertig, 1994): 3He,,; -_ Rtot" 4Hetot - Req" 4He,q - --Ratrn . (20Netot- 20N e,q ) Natm - R,,,. (4He101-4He,q - _1_. (20Ne101_20Ne,q )j Nairn (5.8) 4 4 = He,01 ·(R101 -R,,,)- He,q ·(Req -R,,,)- _1_. (20Netot_20Ne,q). (Ratrn - R,,, ). Nairn In Eq. (5.8), the tritiogenic 3He component is expressed as a function of the measured 4He and 20Ne concentrations and the isotopic ratios of the different components. Atmospheric saturation concentrations are calculated according to Weiss (1971). The helium isotope ratio of the terrigenic component ( R1,,) is assumed to be equal to that of the mantle ( Rman) or that of the crustal component ( Rcr). In some cases it can be determined directly from related samples with a large terrigenic component, e.g. from hydrothermal springs (see Sect. 6.3.3). The 3He/4He ratios of the different components are listed in Tab. 5.1. 83 The square of the error calculated from Eq. (5.8) is (Aeschbach-Hertig, 1994 ): (o3Helri r= 7( a!:lri. &; r 2 2 =(Riot Rrer ) • 04H:ei1 + (R,,q - R,,,, ) · o4H:e;q I II 2 1 2 2 4 +(--·(Ratm - R1,,)) · {0 °Ne;o1 +8 °Ne;q)+ He701 ·OR~r Natm III IV v VI +-2-·1 (20 Netot-· 20 Neeq )2 . (Ratm -Rter )2 ·w•atm.,.,2 Natm VII (5.9) 2 1 20 20 -·{ Ne101 - Ne,q)) ·OR;tm +(-Natm VIII 2 where o denotes the 10' error. This total error is composed of the 4ffe101 (I), 0Ne101 (ill) and R101 (IV) measurement errors, of the errors in the equilibrium values 4ffeeq (II), Req (V) and 20Ne,q (Ill), and of the uncertainties associated with the isotopic ratios Ratm (VII), R1,, (VI) and Na1m (VIII) of the different components. An example to illustrate the relative importance of these different contributions is given in Tab. 5.2. The large uncertainty in the terrigenic helium isotope ratio R1er (.., 80%, see Ch. 6.3.3) is the main contribution to the total error. Note however that the contribution of this error is only important in samples with a large terrigenic component. Other significant sources of error are the measurements of 4He101 (I), 20Ne1a1 (III) and R101 (IV). In the case of no atmospheric helium excess, the error in 20Ne101 does not contribute to the error in 3ffe1n. Uncertainties in the equilibrium concentrations 4He,q (II) and 20Neeq (Ill), reflecting the uncertainty in the water temperature during gas exchange, are responsible for relatively small contributions to the total error. The contributions of the errors in the isotopic ratios of the different components are insignificant. 84 3 Table 5.2: Calculation of 6 He1ri using Eq. (5.9) for a sample taken at the deepest part of the Central Basin of Lake Baikal on June 14, 1992. The Roman numericals correspond to those in Eq. (5.9). 194 0.22 2.76 1.76 0.00 8.24 0.00 0.00 measured quantities: 4Heror =(5.225 ± 0.032)· 10-8 [cm3STP g-11 20Nero1 =(1.908 ± 0.014)· 10-7 [cm3STP g-1 I Rtot =(2.190 ± 0.008)' 10-6 [·] equilibrium values: 4Heeq = (4.548 ± O.Gl4)· l o-8 [cm3STP g-11 (Weiss, 1971) 2~eeq = (1.854 ± 0.006)· IO-7 [cm3STP g-11 (Weiss, 1971) Req (1.359 ± 0.0004)"10-6 I - I (Benson and Krause, 1980) isotopic ratios: Ra.tm =(1.384 ± 0.001)·10-6 [-) (Clarke et al. 1976) Natm = 3.14 ±0.01 [.I (Ozima and Podosek, 1983) Rt er = (2.22 ± l.79)· l o- 7 I -I 85 5.3 Experimental Aspects and Performance Between 1992 and 1995, a total of 281 samples were collected from Lake Baikal for noble gas analysis. The majority of the samples were taken during four expeditions (7 June - 15 July 1992; 18 May - 26 June 1993; 21 October - 14 November 1994). Additional samples were collected by colleagues from the Limnological Institute of the Siberian Division of the Russian Academy of Sciences in Irkutsk. In the laboratory, light noble gases were analysed in different measuring runs according to the procedure described by Kipfer et al. (1994) and Aeschbach-Hertig (1994). An overview of the measuring runs and the difficulties that occurred is given in Tab. 5.3. In the following, a brief outline of the experimental procedure is given together with an account of the measurement performance achieved. Table 5.3: Overview of the samples from Lake Baik.al collected between 1992 and 1995 and the various e.roblems that occurred durin~ the measurements. Year 1992 1993 1994 1995 Samples collected 53 84 69 75 Errors in saml!ling deeth 3 3 6 He-Ne run WA25 WA31 WA'36 WA38 Samples measured 53 84 69 74 Broken Pumped out Air contontaminated 3 7 8 4 Incomplete extraction 3 3 Insufficient sel!aration of He and Ne 2 3H run WA27 WA34 WA37 WA39 Samples measured 52 82 67 59 Broken 2 Pumped out 2 Air contaminated 2 6 7 4 Inefficient extraction 3 2 Unidentified exl!!!rimental £roblems 3 4 86 5.3.1 Sampling and Measurement Procedure Water samples for noble gas analysis were collected in 5-litre Niskin bottles using a winch with a mechanical depth display. The sampling depth error was estimated at 3%. Once on board, the water was immediately transferred from the bottle to specially designed sample containers through a short piece of silicon tubing. Each sample container consists of a 45-ml copper tube mounted on a sliding aluminium rack. The water sample is sealed off tightly by clamping off both ends of the copper tube with stainless steel clamps (Kipfer, 1991 ). The samples were analysed in the laboratory for light noble gases and tritium according to standard institute procedure, which is as follows: The copper tube is connected directly to a UHV system using a modified CF connector and the head space is pumped out to less than 1Q-5 Torr. The clamp on tbe UHV-system side is then loosened carefully and the sample transferred to a 500 cm3 extraction vessel, which is then shaken for 5 minutes to extract the dissolved gases. From the extraction vessel, the gases are transferred to the purification line. Back-diffusion is inhibited by leading the gas flow through a thin capillary to increase gas flow speed. In the purification line, all gases except helium and neon are removed in a series of getters and liquid-nitrogen-cooled traps. Water vapour is removed in a first cold trap. Other atmospheric gases are eliminated in two nitrogen-cooled zeolith traps. All remaining reactive substances are caught in two getters (NP 10; Zr-Ti bulk getter) and nitrogen-cooled active charcoal. Finally, helium and neon are separated by freezing neon to a cryogenic cold trap at 14 K. Both helium isotopes are then analysed simultaneously in a non-commercial double- collector 90° magnetic sector spectrometer (r= 210 mm). The spectrometer is equipped with a Baur-Signer source (GS 98) tuned for maximum linearity (Baur, 1980). The 4He current is integrated in a Faraday cup. 3He is detected and counted in a 16-stage, discrete Cu-Be dynode electron multiplier. The mass resolution (defined as the average mass m devided by the mass difference Om of two neighbouring ions) of about 750 allows accurate separation of 3He+ and HD+ ion signals (Dunai and Touret, 1993). As soon as the helium measurement has started, the neon is released from the cryogenic cold trap and analysed in sequence in a second non-commercial 90° magnetic sector spectrometer (r 120 mm) equipped with a Baur-Signer source (Baur, 1980). 20Ne and 22Ne are registered using a Faraday cup. After noble gas analysis, the degassed water sample is transferred back into the original copper tube and tightly sealed again. Several months later, the sample is reanalysed in a simplified process for the 3He produced by the decay of tritium during storage. This procedure allows the subsequent determination of light noble gases and tritium concentrations from the same sample. 87 S.3.2 Calibration Noble Gases Noble gas concentrations and isotopic compositions were calculated from the measured signals by calibrating against an air standard (Aeschbach-Hertig, 1994). To compensate for short-term sensitivity changes of the spectrometer, fast calibrations (fastcals) were carried out before each noble gas analysis. The fastcal gas, which is a pure helium or neon standard, was directly injected into the spectrometer. The so-called fastcal amount FA;, which relates the calibration gas to the sensitivity changes of the spectrometer, is calculated as (5.10) The fascal amount [cm3STP] is a virtual quantity combining the volume of the calibration gas CA; [cm3STP] with the mean signal of the fastcals measured before and after each calibration i, FSi, and the calibration signal CS; (3He signal in [cps]; 4He and 20Ne signals in [A]). The gas volume dissolved in a sample j, SAj [cm3STP), is calculated as (5.11) where SSi is the measured signal of sample j, FS j the mean signal of the fastcals measured before and after samplej and FA is the long-term mean fascal amount [cm3STPJ. Dividing SAj by the sample weight yields the noble gas concentration in cm3STP g-1. The error in the 3ffe signal (approximately 0.4%) is due mainly to the counting statistics. The internal errors in the 4ffe and 20Ne signals (0.03% and 0.1 - 0.3%, respectively) are generally much smaller and are mainly a result of the uncertainty associated with the linear regression applied to compensate for ion-pumping during the analysis. Compared to the long-term stability of the measurements, these errors are negligible. The total error in the measurement, including that associated with gas extraction and purification, corresponds basically to the reproducibility of the measurement as expressed by the standard deviation of the fastcal amounts (Aeschbach-Hertig, 1994). For quality control reasons, aliquots of an internal freshwater standard are routinely measured. The standard deviations of the measured concentrations of these aliquots agree fairly well with those of the fastcal amounts. 88 Tritium Tritium is measured indirectly by reanalysing the degassed sample after several months for the 3He produced by the decay of tritium. Since the analysed gas consists practically of pure 3He, it is calibrated only against the fastcal helium standard. The volume of 3He1ri in the sample is calculated by multiplying the measured signal by the mean ratio of the volume of the calibration gas to the calibration signal. Compared to the noble gas analysis, where the 3He signal is on the order of 50 cps, the 3He signal of the tritium analysis is much smaller ("' 1 cps). The error in the counting statistics is therefore much larger("' 3% ). 5.3.3 Performance 5.3.3.1 Errors in Sampling Depth 3ffe concentrations in samples collected at the deepest part of the Southern Basin (position Sl, see Fig. 6.1) between 1992 and 1994 are shown in Fig. 5.2. In 1993, the concentrations at 600 m, 1000 m and 1200 m depth were significantly lower than in 1992 and 1994. Noble gas concentrations measured in samples that were originally collected for freon analysis confirm these results. Thus, experimental problems can be excluded as a cause of the observed deviations. On the other hand, the low 3He concentrations in 1993 are not the result of a major deep-water renewal event. Between 1993 and 1994, 3He concentrations increased and reached values comparable to those observed in 1992. However, the tritium concentrations in Lake Baikal (Fig. 6.4) were too low to explain the observed increase in 3ffe by radioactive decay. Within the time between the expeditions (l.5 yr), the radioactive decay of tritium would cause the 3He concentrations to increase only by about 0.3· l o-14 cm3STP g-1. We therefore concluded that the low values in 1993 were caused by an error in sampling depth larger than 3%. Most likely the sampling bottles closed too early during their descent, so that the samples actually originate from above 600 m depth. The above example illustrates that incorrect sampling depths can only be identified by comparing the measured concentrations with reliable results obtained from samples collected at the same location. Since it is not possible to correct for this effect reasonably, no samples identified as originating from an incorrect sampling depth are included in the discussion in Sect. 6. This problem occurred most frequently during the autumn 1994 expedition (Tab. 5.3) when limited control over sampling depth was experienced due to severe storms. During 89 the spring expeditions in 1992, 1993 and 1995, winds were generally weak and less samples were identified as originating from incorrect sampling depths. 0 c. b. b.b. 500 ] 0 • b. ~ =QI b. Q 1000 Figure 5.2: 3He concentration at the deepest part • of the Southern Basin of Lake Baikal in 1992 (b.), 1993 (.)and 1994 (0). In 1993, the samples from • 600 m, 1000 m and 1200 m depth were most likely collected from the wrong depths. 1500 6 7 8 9 10 11 3 14 3 1 He [10" cm STP g" ] 5.3.3.2 Noble Gas Analysis Relative Neon Saturation: Neon, analysed simultaneously with helium, served mainly as an indicator of the degree of atmospheric contamination of a sample. Since the atmosphere is considered to be the only source of dissolved neon, an atmospheric helium excess can be identified by a neon excess. In the majority of the 281 samples, measured neon concentrations deviated by less than 5% from the atmospheric equilibrium value at in situ temperature. A neon excess exceeding 5% was observed in IO samples (5 samples taken from hydrothermal springs and 5 samples contaminated with air). Relative under-saturation by more than 5% was observed in 4 samples. Obviously, the gases dissolved in these samples were extracted incompletely. The relative neon saturation in the remaining 267 samples is shown in Fig. 5.3. The data are distributed around a mean value of x = 0.42% and can be approximated well with a Gaussian distribution (cr = 0.63). The positive mean value implies that there is a significant neon excess, which is most likely caused by the injection of air bubbles through the water surface. The 3cr confidence interval is given by (0.42 ± l.89)%. Deviations from the atmospheric equilibrium value of more than ± 2% are therefore considered as indicative of excess air or of incomplete extraction of the sample. Neon excesses exceeding 2% were observed in 12 samples. In 3 samples, neon was under-saturated by more than 2%. All 90 Figure 5.3: Distribution of 20 the relative neon saturation anomaly. The data are distributed around a mean 15 of x = 0.42% and can be !l c:: approximated well with a Gaussian distribution. 22 Q= 10 u samples have a neon excess larger than 2%. 7 samples 5 are neon under saturated by more than 2%. 0 .4 -2 0 2 4 A2lNe [%] samples with &Ne :> 2% were corrected for air for the total neon excess. Samples with &Ne < - 2% were not considered in the discussion in Sect. 6. Fas teal Amounts and Freshwater Standards Mean fastcal amounts of 3He, 4He and 20Ne calculated from Eq. (5.10) are listed in Tab. 5.4. In the four measuring runs, the mean FAs of lHe and 4He were fairly constant. Only in the case of neon did the mean FA vary considerably between different runs, mainly because of changes in the sensitivity of the spectrometer. Between W A25 and W A3 l, the mean neon FA increased by more than a factor of 3 due to the use of a new neon calibration standard. The standard deviations of the values illustrate the reproducibility of the measurements. For all isotopes this is clearly better than I%. Table 5.4: Meanfastcal amounts (x) of3He, 4 He and 20Ne. The standard deviations (a) represent the reproduceability reached in the different runs. It is clearly better than !%. lHe [10-I l cml STPJ 4He [lo-5 cm3 STPJ 2~e [J0-5 cm3 STP] Sampling Run x 1992 WA25 0.83188 0.00577 0.56213 0.00289 0.19215 O.OO!lO i0.69%) i0.51%) i0.57%l 1993 WA31 0.82819 0.00380 0.56076 0.00206 0.71258 0.00347 (0.46%) (0.37%) (0.49%) 1994 WA36 0.82624 0.00738 0.56036 0.00436 0.71395 0.00358 (0.89%) (0.78%) (0.50%) 1995 WA38 0.81544 0.00702 0.55321 0.00422 0.67510 0.00613 (0.86%) (0.76%) (0.91%l 91 .. 0.05 0.05 ..!. '"". -.~ o.o "' .c"' .c. -0.05 Figure 5.4: Helium and neon concentrations measured in -. aliquots of an internal freshwater standard normalised to the respective mean concentrations. The aliquots were collected on February 13, 1992. Samples from Lake Baikal were measured in runs WA25, WA31, WA36 and WA38. 0 500 1000 1500 Time since sampling [d] Aliquots of an internal freshwater standard were routinely measured to monitor the reproducibility of the helium and neon measurements, including gas extraction, purification and gas separation. Helium and neon concentrations of these aliquots are shown in Fig. 5.4. The measured concentrations are distributed around well-defined means. Standard deviations are less than 0.7% of the mean concentrations for 4He and 20Ne, and about 0.9% for 3He. 92 S.3.3.3 Tritium Measurements 4He Concentration aml Blank: The 4He concentration analysed simultaneously with the tritiogenic 3ffe served as an indicator for air contamination during storage or during the measurement, and for insufficient extraction when analysing the sample for noble gases. The mean 4He concentrations in the samples from Lake Baikal measured in each nm are listed in Tab. 5.5. These mean values are close to zero, and standard deviations are on the order of 10-10 cm3STP g-1. Deviations from the 3a interval are most likely caused by insufficient extraction or air contamination. 4He concentrations larger than 2·10-9 cm3STP g-1 are considered as indicative of air contamination (Aeschbach-Hertig, 1994). In WA37, air contamination occurred in about 10% of the samples (Tab. 5.3). The large contamination rate was most probably caused by a leaking valve that was not detected before WA39. In WA27, WA34 and WA39, the contamination rate was significantly smaller. Table S.S: Mean 4He concentrations measured during the tritium analysis. 4He [J0-10 cm3 STP] Sampling Run 1992 WA27 -0.03 0.57 1993 WA34 0.50 1.20 1994 WA37 -l.15 2.83 1995 WA39 0.61 2.42 Immediately after the tritium measurement, the samples were reanalysed for remaining 3ffe to monitor the efficiency of the extraction. The signals of these blank measurements (Fig. 5.5) are distributed around a mean value of 0.009 cps and can be approximated by a Gaussian distribution ( <1 = 0.013). Blank signals larger than 0.045 cps are considered as indicative of an insufficient first extraction. This was the case for 12 of the 243 blank measurements. Standards 3ff concentrations of the internal freshwater standards are shown in Fig. 5.6. The 27 aliquots measured before W A37 yield a mean concentration of 30.12 ± 0.95 TU. The concentration of the aliquot measured in WA39 is in good agreement with the long-term average. Compared to 93 35 30 25 ~ :I 20 Figure 5.S: Distribution of 0 u 3 He signals in blank 15 measurements following the 10 tritium analysis. Signals larger than 0.045 cps are 5 considered as indicative of an insufficient first extraction. -0.05 0 0.05 0.1 3He signal [cps] this, the concentrations of the aliquots measured during W A37 are unreasonably large; they exceed the mean concentration by up to 25%. Furthermore, the values vary considerably. The reason for the large deviation of the concentrations measured during W A37 from the long-term mean value remains unclear. Most likely, the high values were caused by the so-called dark count rate, i.e. by self-induced counts of the multiplier. The dark count rate may cause a problem especially for samples with low 3He concentrations. In this case, it remains unclear why the blank measurements, where the 3He concentrations are practically zero, were hardly affected by this phenomenon. As indicated by the considerable variation of the concentrations of the aliquots, not all samples were affected to the same extent, and thus it is not possible to reasonably correct for this effect. Figure 5.6: Tritium concentrations measured in aliquots of an internal fresh- water standard normalised to the respective mean concen- trations. The aliquots were collected on February 13, 1992. Samples from Lake Baikal were measured in runs WA27, WA34, WA37 and WA39. 0 500 1000 1500 Time since sampling [d] 94 S.3.4 Overall Performance From the 281 samples collected, 244 measurements yielded reliable helium concentrations (Tab. 5.3). Air contamination was observed in 22 samples; an additional 15 samples were lost completely for various reasons during noble gas analysis. From the 260 samples that were analysed for tritium, a total of 35 were lost. The results from samples collected in 1994 tended to be less reliable than those from samples collected in other years, mainly for two reasons. (l) Because of strong winds we had limited control over the sampling depth. At least 7 samples were identified as being taken from incorrect depths. (2) The tritium concentrations of aliquots of an internal freshwater standard sometimes deviated significantly from the long-term mean value. The reason for these large deviations remains unclear. Since both effects cannot be reasonably corrected for, the data from 1994 were not included in the discussion in Sect. 6. The overall reliability of the samples collected in 1992, 1993 and 1995 was satisfactory. This is especially true of the samples collected in 1992 and 1995. These data will be discussed in the following section. 6. Distribution of Helium and Tritium in Lake Baikal R. Hohmannll, M. Hoferll, R. KipferO, F. Peeters]), H. Baur2l, M. N. Shimaraev3l, and D. M. lmbodenll llswiss Federal Institute of Technology (ETII) and Swiss Federal Institute of Environmental Science and Technology (EAWAG), 8600 Dilbendorf, Switzerland 2lJ.sotope Geology, Swiss Federal Institute of Technology, NO C61, 8092 Zilrich, Switzerland 3Jumnological Institute of the Siberian Division of the Russian Academy of Sciences, Irkutsk 664033, Russia (accepted by Journal ofGeoghysical Research) Abstract The 3H-3He age of a water mass is a measure of the time that has passed since the water mass was last in contact with the atmosphere. Between 1992 and 1995, a detailed study of 3H-3He ages was conducted in Lake Baikal, the deepest and largest lake by volume on Earth, to investigate deep-water renewal in its three major basins. Maximum 3H-3He ages are 14 - 17 yr in the Southern Basin, 16 - 18 yr in the Central Basin and IO - 11 yr in the Northern Basin. Rates of renewal of deep water with surface water, deduced from volume-weighted mean 3H-3He ages below 250 m depth, are about IO% yrl in the Southern and Central Basins and 15 % yr-1 in the Northern Basin. In the Southern Basin, the mean 3H-3He age below 250 m depth increased steadily from 9.6 yr in 1992 to 11 yr in 1995, indicating a slight diminution in deep-water renewal during this time. Bottom-water renewal by large-scale advection was estimated from the mass balance of 3He in the 200 m thick bottom layer of each basin. Bottom- water renewal rates in the Northern Basin were found to be between 80 and 150 km3 yrl and in the Central Basin between 10 and 20 km3 yrl, whereas in the Southern Basin they were practically zero. Correlating oxygen and dissolved helium-4 concentrations with 3H-3He age allowed us to determine the mean hypolimnetic oxygen depletion rate in the water column (4.5 µmol 1-1 ycl), as well as mean helium fluxes from the lake bottom (2.8·1011 atoms m-2 s-1 in the Northern Basin, and 1.3· l Oil atoms m-2 s-1 in the Central and Southern Basins). The helium isotope ratio of the terrigenic helium component injected from the lake bottom, determined from measurements of water from hydrothermal springs in the vicinity of the lake, was found to be approximately 2.2·10--7. 96 6.1 Introduction Lake Baikal, located in the Great Baikal Rift zone of eastern Siberia, is the deepest ( 1632 m) and largest (23015 km3) lake by volume on Earth. It holds about one fifth of the global fresh (liquid) surface water. The lake is divided by underwater sills into three main basins (Fig. 6.1 ): the Southern Basin (max. depth: 1432 m), the Central Basin (1632 m), and the Northern Basin (897 m). The three major rivers entering the lake, the Selenga, the Upper Angara and the Barguzin, account for more than 70% of the total annual river inflow. The only outflow is the Lower Angara at the south-western end of the lake. Lake Baikal is the oldest of all existing lakes: the Baikal tectonic depression is thought to have formed as early as the Oligocene (Golubev et al., 1993). Basin sediments, which are several kilometres deep, record more than 15 Myr of the lake's history (Members, 1992) During this time, a rich endemic flora and fauna has evolved. Despite the lake's great depth, dissolved oxygen concentrations exceed 80% relative saturation thrughout the entire water column. The high oxygen concentrations indicate ( 1) that oxygen consumption in Lake Baikal is small and (2) that deep-water renewal is surprisingly rapid. The latter point raises the question of the intensity of deep-water renewal and of how it occurs in this deep lake. A first estimate of the deep-water renewal rate was derived by Verbolov and Shimaraev (1973) from the product of estimated vertical velocities and the estimated areas of regions where deep-water formation is likely to occur. The authors concluded that complete exchange between surface and internal waters takes place within 11 yr. First 3H-3He ages were determined by Craig (1989), who reported a maximum 3H-3He age of 18 ± 4 yr for the Central Basin. Weiss et al. (1991) showed that water ages calculated from the vertical distribution of chlorofluorocarbon-12 (CFC-12) are less than 16 yr in the entire lake. Based on these data and on dissolved oxygen concentrations, Killworth et al. (1996) developed a one-dimensional inverse model to calculate annually averaged fluxes. Their "ventilation times" below 400 m depth, derived by fitting modelled tracer concentrations to observed data, are in good agreement with the CFC-12 ages obtained by Weiss et al. ( 1991 ). A three-dimensional model to study the intensity and scale of plumes involved in the formation of deep water in deep temperate lakes was developed by Walker and Watts (1995). Weiss et al. (1991) suggested that the so-called thermobaric instability, which is linked to the pressure dependence of the temperature of maximum density (TmtJ), is the key process controlling deep-water ventilation in Lake Baikal. Shlmaraev et al. ( 1993) interpreted the thermal bar which develops along most of the shoreline of Lake Baikal in spring as the main trigger of deep-water formation. However, Peeters et al. (1996) were able to show that 97 55 Lake Baikal 54 53 52 0 so 100 km 104 105 106 107 108 109 110 Longitude [0 E] Figure 6.1: Map of Lake Baikal with isobaths at 400, 700 and 1000 m depth. The Selenga, Upper Angara, and Barguzin rivers are the main inflows, the Angara River at the south-western end of the lake being the only outflow. The lake is divided into three basins: SJ, Cl and NJ mark the deepest points (e) of the Southern, Central and Northern Basins, respectively. Samples from S2, C2, C3, C4 and N2 (®) were included in the calculation of volume-weighted mean concentrations. Hydrothermal springs (0) are Khakusy Hot Spring (Ksy), Frolikha Bay (FB). Zmeinnyi lstochnik (Z) and Kotelnikovsky (Kot). K2 (Cl) is a sampling station in the Kukui Canyon. 98 convection related to the thermal bar is restricted to the uppermost 500 m of the water column. A more general analysis of processes of deep-water renewal, including the effect of salinity on water density, was presented by Hohmann et al. (1997a) Based on a thorough theoretical analysis of stability and neutrally buoyant transport in lakes (Peeters et al., 1996) and on a detailed set of CTD data, the authors showed ( 1) that the Selenga River is important for deep-water renewal in the Central Basin and most likely also in the Southern Basin, and (2) that deep-water renewal in the Northern Basin is brought about mainly by a combination of small salinity and temperature gradients between the Central and Northern Basins. The 3H-3He method was introduced into oceanographic research by Jenkins and Clarke (1976) and was subsequently applied in numerous studies (e.g., Andrle et al., 1988; Jenkins, 1987; Schlosser et al., 1990). Compared to its use in oceanography, relatively little effort has been made to employ it routinely in limnology, despite the fact that the time-scale covered by the 3H-3He method matches the time-scale of mixing processes in lakes (Schlosser, 1992). Torgersen et al. (1977) were the first to apply this method to limnic systems, using it to estimate gas exchange rates, gas renewal at turnover and vertical diffusivity in the epilimnion in Lakes Erie, Huron and Ontario. Since then, the method has been applied successfully in the analysis of vertical exchange, exchange between lake basins, oxygen consumption and gas exchange, groundwater influx and helium flux from the lake bottom (e.g. in Lake Baldegg (Switzerland) (Imboden et al., 1981), Green Lake (NY) (Torgersen et al., 1981), Konigsee (Germany) (Fischer et al., 1985), Lake Constance (Germany) (Zenger et al., 1990), Lake Van (Turkey) (Kipfer et al., 1994) and Lake Lucerne (Switzerland) (Aeschbach- Hertig et al., l 996a)). In this paper we present a detailed set of helium, neon and tritium concentrations from samples collected in Lake Baikal between 1992 and 1995. Deep-water renewal rates in the three major basins are derived from 3H-3He ages calculated from 3H and 3He concentrations. Oxygen consumption rates and the 4He flux from the lake bottom are estimated by correlating observed values of the oxygen deficit and 4He excess with calculated 3H-3He ages. Advective bottom-water renewal is calculated from the 3He mass balance in the deepest 200 m layer of each individual basin. A six-box model of Lake Baikal which bases on the data presented here and on CFC-11 and CFC-12 concentrations is discussed by Peeters et al. (1997). 99 6.2 Methods 6.2.1 CTD Measurements Temperature, conductivity and pressure were recorded in situ with an SBE-9 on-line CTD probe from Seabird Electronics. The probe has a resolution of 0.025 dbar for pressure, 0.0003 K for temperature and O.Ql µS cm-1 for salinity. The usual drift of the temperature sensor is 0.004 K yrl. Oxygen concentrations were calibrated with Winkler titration data (accuracy: 0.3%) provided by T. Khodzher (unpublished data). Ionic salinity Sc, defined as the total mass of dissolved ions per unit mass of solute, was detennined from the electrical conductivity and average chemical composition of Baikal water (Falkner et al., 1991) according to the procedure described by Hohmann et al. (1997a). Potential temperature 8 and water density were calculated from the CTD data and the concentration of silicic acid (Si( OH)4), the only non-ionic species that significantly influences spatial density gradients. Due to the lack of more detailed information, the distribution of silicic acid for the period 1994/95 was determined from unpublished data taken by T. Khodzher in 1993. The quasi-density Pqua• a measure which proves to be most adequate for detennining the local stability of the water column in deep lakes, was calculated according to Peeters et al. (1996) as follows: Z, Pqua(z,z,) = p(z)- J'l'(z' ,i )dz', (6.1) z where 'l'(z' ,z' )is the local adiabatic density lapse rate of the background field at depth z': (6.2) The quasi-density depends on the distribution of 8, S, and p in the water column. The gradient of Pqua is proportional to the Brunt-VaisiiUi frequency N1. A water column is locally stable if Pqua increases with depth (i.e. if Pqua decreases with z). 6.2.2 Noble Gas Analysis Water samples for noble gas analysis were taken using 5-1 Niskin bottles. The error in the sampling depth was estimated as ± 3%. On board, the water was immediately transferred to 45-ml copper tubes which were then tightly sealed using pinch-off clamps. 100 In the laboratory, light noble gases were analysed according to the procedure described by Kipfer et al. (1994). The samples were connected to the inlet of a UHV system and the dissolved gases extracted from the water were transferred into the purification section, where all gases except He and Ne are removed by several liquid-nitrogen-cooled traps and getters (S.A.E.S. C50, NP 10, Zr-Ti bulk getter). The He and Ne remaining was then separated in a cryogenic cold trap and the concentrations measured in two different static mass spectrometers. After extraction, the degassed water was transferred back into the original copper tube. Several months later, the samples were reanalysed for 3He produced by the decay of tritium. Noble gas concentrations and isotopic composition were regularly calibrated against an air standard. Aliquots of an internal freshwater standard were routinely analysed for quality control purposes. The standard deviation of the measured concentrations of these aliquots corresponds to the reproducibility of the measurement procedure, including that associated with gas extraction and purification. Except for the samples collected in fall 1994, the reproducibility was better than± 0.5% for the helium isotope ratio and about± 1% for noble gas concentrations. The tritium concentrations calculated from the 3He produced by the decay of tritium during storage have an error of± 3%. To account for the additional uncertainty related to sampling depth, the estimated depth error (± 3% of total depth) was multiplied by the absolute value of the mean gradient of the measured concentrations. The total uncertainty in the concentrations and ratios at a certain depth was calculated as the square root of the sum of the squares of the individual errors. The quality of the samples collected in 1994 turned out to be impaired for two reasons: (1) because of stormy weather the sampling depth was less precisely known, and (2) a leakage in the purification line affected some of the samples. Since it was not possible to eliminate these effects, the 1994 data will not be considered in the following discussion. 6.2.3 Calculation of 3H.3He Age The 3H-3He age 'twas calculated from the following basic equation (Torgersen et al., 1977): (6.3) where A.= 0.05576 yr I is the decay constant of tritium (Unterweger et al., 1980). 3H and 3Hetri are the concentrations of tritium and tritiogenic helium, i.e. the helium produced by the decay of tritium. The former concentration is usually given in TU (1 TU being equivalent to a tritium/hydrogen ratio of 1()-18), the latter in cm3 STP g-1 (1 cm3STP g-1 =4.019· 1014 TU). IOI The tritiogenic helium concentration 3He1,; was calculated as follows: 3H etri-_3H em-3H eeq-3H eex-3H eter (6.4) 4 4 4 4 =Rm· Hem Req· Heeq -Rex· Heex -R1,,- He1er, where the subscript m refers to measured concentrations, eq to solubility equilibrium concentrations, ex to atmospheric excess concentrations, ter to terrigenic concentrations, and R refers to the corresponding 3He/4He ratios. Equilibrium concentrations were calculated according to Weiss (1971). The helium isotope ratio in water in equilibrium with the atmosphere is Req = 1.36· 1()-6 (Benson and Krause, 1980). The atmospheric helium excess 4Heex is assumed to originate from the complete dissolution of air (e.g. from injected air bubbles) without fractionation. It was calculated only for samples with a neon excess (defined as the difference between the measured Ne concentration and the atmospheric equilibrium concentration) greater than 2 %, which is indicative of atmospheric contamination. For such samples, 4He,x was calculated by multiplying the total Ne excess with the atmospheric ratio 4He/20Ne (0.318; Ozima and Podosek, 1983). The atmospheric helium isotope ratio is Ratm = 1.384· I0--6 (Clarke et al., 4 1976). 4He1,, is the He excess remaining after correcting for air. It comprises helium from both the Earth's mantle and crust. The value of R1er is discussed below. More details regarding the noble gas analysis are given by Kipfer et al. (1994) and Aeschbach-Hertig et al. ( 1996b ). A detailed description of the calculation of the 3H-3He age can be found in Aeschbach-Hertig (1994) and Aeschbach-Hertig et al. (1996a). 6.3 Results The data presented here were collected during four cruises organised by the Baikal International Center of Ecological Research (BICER) during 5 - 15 July 1992, 18 May - 26 June 1993, 21 October 14 November 1994 and ll May -4 June 1995. From aboard the research vessel RV Vereshchagin, more than 600 CTD profiles were recorded and a total of 281 water samples for noble gas and tritium analysis were collected. 6.3.1 Vertical Density Stratification In Fig. 6.2, profiles of potential temperature (J, ionic salinity Sc, oxygen concentration [02], and quasi-density Pqua (Peeters et al., 1996) are illustrated. These profiles were measured at the deepest station of each basin (marked by• in Fig. 6.1) in May 1995, shortly after break- 102 0 200 N -- -\ 400 \ 600 ~ ] N' 800 =1000 Q2' 1200 1400. ...-a.> c 1600 1.5 2 2.5 3 3.5 94.5 95.0 95.5 1 0 [°C] Sc[mgkg" ] 0 -~ ...... ; 200 L 400 \ \ \ \s ] 600 ... N' L \ 800 \ I)., ... I =Q 1000 \ 1200 \ 1400 c °\ c) d) 1600 0.055 0.06 0.065 350 400 3 p -1000 [kg m" ] 1 qua (02] [µmo! r 1 Figure 6.2: (a) Potential temperature fJ, (b) ionic salinity Sc- (c) quasi-density Pqua• and (d) dissolved oxygen concentration {02) at the deepest point of the Southern Basin (dotted line), the Central Basin (solid line) and the Northern Basin (dashed line) measured from May 18 24. 1995. The straight dotted line in (a) is the temperature of maximum density, Tmd· The influence of non-ionic silicic acid was included in the calculation of quasi-density. up. The temperature profiles (Fig. 6.2a) are characteristic of a typical winter stratification. Near the surface, the water column is inversely stratified with respect to temperature. The temperature reaches its maximum, the so-called mesothermal temperature maximum, where it is equal to the temperature of maximum density (thin dotted line in Fig. 6.2a) at about 200 m. Below this depth it decreases steadily. The bottom temperature is about 3.4 °C in the Southern and Northern Basins and 3.1 °C in the Central Basin. Ionic salinity is low and its gradients (Fig. 6.2b) are rather small. Below 50 m, and between the individual basins, salinity varies by less than I mg kg-1. The vertical salinity 103 distribution in the Southern and Central Basins is similar. The destabilising effect of the slight salinity drop with depth in the interior of these basins is compensated for by the gradient of 8. In the Central Basin, there is a distinct salinity increase in the deepest 100 m that is also observed in the Southern Basin, although less pronounced. In the Northern Basin, salinity is significantly lower than in the other basins. Here, it increases steadily with depth and thus contributes to the vertical stability of the water column. The maximum value close to the surface in the Northern Basin is most likely due to the salinity increase resulting from the formation of ice in early winter. The steady increase in quasi density with depth in all three basins (Fig. 6.2c) illustrates that the combined effects of temperature and salinity (including the contribution of silicic acid) is to produce a stable stratification (Peeters et al., 1996). The thermocline at about 100 m depth marks the zone of maximum stability, i.e. the zone with the largest gradient of pqua. This gradient is small near the mesothermal temperature maximum indicating that the stratification is weak. A pronounced increase in Pqua near the bottom of the Central Basin implies the existence of a stably-stratified boundary layer. Despite the great depth of Lake Baikal, oxygen concentrations are high throughout the entire water column (Fig. 6.2d). Even at the bottom of the lake, the relative oxygen saturation generally exceeds 80%. 6.3.2 Distribution of He, Ne and 3ff in the Water Column Table 6.1 summarises the concentrations of helium, neon, and tritium observed between 1992 and 1995 at the deepest station of each basin. In Fig. 6.3, the helium isotope ratio 3He/4He is plotted against the corresponding elemental Ne/He ratio (expressed as 20Nef4He). The 3He/4He ratios of all samples clearly differ from the 3He/4He ratio of air-saturated water (ASW) established by gas exchange at the lake surface. Provided that there is no input of helium from the Earth's mantle, deviations from atmospheric equilibrium toward larger 3He/4He ratios are generally caused by excess 3He produced by the decay of tritium (Torgersen et al., 1977; Imboden et al., 1981; Aeschbach-Hertig et al., 1996a). In contrast, deviations from atmospheric equilibrium toward lower 20Nel4He ratios indicate the presence of a crustal component which is virtually free of neon and has a low 3He/4He ratio (Aeschbach-Hertig et al., l 996a; Kipfer et al., 1996). Because for most 20Ne concentrations (except for the contaminated samples of the 1994 sampling campaign) deviations from atmospheric equilibrium were less than 2%, the observed helium anomalies cannot be explained in terms of an atmospheric source. 104 Table 6.1: Noble gas and tritium concentrations measured at the deepest station of each basin in the years 1992, 1993, and 1995. Errors are given as 1 CJ values. Symbols are used as follows: * 3 H lost due to experimental problems, value calculated from linear regression; + Ne excess > 2%; · Ne lost due to experimental problems. slat date depth T 4He 3ffe/4He 20Ne/4He 3Hen 4eeex 20Neex jH [mJ rcJ [10.j\ 110·6] [JO•l4 (10-9 (10-9 rrui cm3STP [IJ cm3STP s·lj cm3STP g·lj cm3STP g·1J SI 23.3.92 0 0.5 4.67±0.03 1.390±0.008 4.23±0.04 0.21±0.05 0.52±0.29 6.74±1.56+ 21.6±0.6 SI 23.3.92 so I 4.70±0.03 I .455:t0.008 4.02±0.04 0.58±0.06 0.88±0.29 -1.13±1.48 19.4±0.5 SI 26.3.92 150 3 4.93±0.03 1.569±0.008 3.98±0.04 1.55±0.06 3.82±0.30 10.12±1.46+ 18.6±0.5 SI 26.3.92 300 3.5 4.66±0.03 l.656±0.011 4.01±0.04 1.54±0.07 1.18±0.29 1.55±1.45 18.4±0.6 SI 26.3.92 400 3.5 4.62±0.03 1.719±0.013 4.01±0.04 1.76±0.08 0.75±0.28 0.18±1.47 18.3±0.6 SI 26.3.92 460 3.5 4.64±0.03 l.782:t0.014 4.00±0.04 2.10±0.08 1.00±0.29 0.43±1.48 18.3±0.5 SI 26.3.92 700 3.4 4.69±0.03 1.986±0.018 3.99±0.04 3.15±0.10 1.50±0.29 2.16±1.55 16.7±0.5 SI 26.3.92 900 3.4 4.69±0.03 2.139±0.022 3.97±0.04 3.86±0.12 l.51±0.30 1.19±1.44 16.1±0.5 SI 26.3.92 1100 3.4 4.71±0.03 2.329±0.025 3.94±0.04 4.77±0.14 1.66±0.30 0.35±1.40 16.0±0.6. SI 26.3.92 1350 3.4 4.76±0.03 2.220±0.030 3.92±0.04 4.40±0.16 2.23±0.32 1.64±1.49 15.9±0.4 SI 19.5.93 0 2.6 4.63±0.02 1.436±0.008 4.06±0.03 0.46±0.05 0.74±0.18 1.29±1.34 18.8±0.5 SI 19.5.93 1000 3.4 4.85±0.02 2.197±0.024 3.83±0.03 4.49±0.12 3.12±0.20 0.68±1.30 15.3±1.3 SI 19.5.93 1200 3.35 4.95±0.02 2.283±0.027 3.76±0.03 5.12±0.14 4.10±0.22 1.12±1.22 14.8±0,4 SI 19.5.93 1430 3.35 4.90±0.02 2.295±0.032 3.81±0.03 5.06±0.16 3.59±0.23 1.50±1.41 14.9±0.5 SI 14.5.95 20 1.5 4.73±0.04 1.400±0.008 4.06±0.05 0.38±0.07 1.37±0.38 2.95±1.79 19.2±1.2 SI 14.5.95 200 3.55 4.68±0.04 1.597±0.010 3.99±0.05 1.30±0.08 1.41±0.38 l.57±1.57 17.0±0.7 SI 14.5.95 600 3.5 4.80±0.04 l.846±0.017 3.87±0.05 2.69±0.11 2.61±0.39 0.99±1.56 15.8±0.7 SI 14.5.95 1000 3.4 4.70±0.04 2.268±0.024 4.09±0.05 4.48±0.15 1.56±0.39 6.83±1.56 14.3±0.7 SI 14.5.95 1200 3.35 4.86±0.04 2.316±0.028 3.82±0.04 5.09±0.17 3.24±0.41 0.80±1.52 13.4±0.7 SI 14.5.95 1300 3.35 5.00±0.04 2.316±0.029 3.70±0.04 5.40±0.17 4.60±0.42 -0.19±1.54 14.9±0.7 SI 14.5.95 1429 3.4 4.79±0.04 2.355±0.032 3.89±0.04 5.11±0.18 2.49±0.41 1.09±1.54 12.9±0.6 Ml I4.6.92 0 3.5 4.70±0.03 1.444±0.008 3.99±0.04 0.63±0.06 l.65±0.29 2.90±1.43 19.3±0.6 Ml 14.6.92 100 3.5 4.65±0.03 l.495±0.008 4.04±0.04 0.79±0.06 1.14±0.28 3.22±1.44 20.3±0.5 Ml 14.6.92 300 3.5 4.66±0.03 1.623±0.011 3.97±0.04 1.40±0.07 1.24±0.29 0.20±1.59 20.2±0.6 Ml 14.6.92 500 3.4 4.68±0.03 l.740±0.026 3.98±0.04 2.06±0.13 1.45±0.29 1.34±1.54 19.3±0.6* Ml 14.6.92 930 3.3 4.81±0.03 2.259±0.023 3.89±0.04 4.70±0.13 2.64±0.31 2.12±1.46 17.2±1.5 Ml 14.6.92 1200 3.2 4.77±0.03 2.488±0.028 3.91±0.04 5.69±0.16 2.22±0.31 l.50±1.58 15.2±0.5 Ml 14.6.92 1400 3.2 4.84±0.03 2.559±0.032 3.85±0.04 6.21±0.18 2.91±0.32 1.59±1.37 14.1±0.5 Ml 14.6.92 1620 3.2 5.23±0.04 2.190±0.036 3.65±0.04 5.27±0.19 6.75±0.35 5.85±1.42+ 15.2±0.5 Ml 31.5.93 200 3.6 4.63±0.02 1.573±0.009 3.99±0.03 l.11±0.05 0.87±0.18 0.29±1.29 18.8±0.4 Ml 31.5.93 600 3.4 4.67±0.02 l.824±0.015 3.96±0.03 2.35±0.08 1.30±0.18 0.07±1.21 18.7±0.5 Ml 31.5.93 953 3.3 4.70±0.02 2.419±0.023 3.94±0.03 5.20±0.12 1.54±0.20 0.49±1.31 15.1±1.2 Ml 31.5.93 1334 3.2 4.72±0.02 2.491±0.030 3.94±0.03 5.57±0.15 l.69±0.22 1.02±1.16 14.3±0.4 Ml 31.5.93 1500 3.2 4.70±0.02 2.502±0.034 3.97±0.03 5.57±0.17 1.47±0.22 l.36±1.35 14.3±1.2 Ml 31.5.93 1620 3.1 4.72±0.02 2.444±0.035 3.91±0.03 5.34±0.18 l.66±0.23 ·l.55±1.18 13.8±0.4 Ml 16.5.95 20 1.2 4.70±0.04 1.377±0.008 4.03±0.05 0.21±0.07 0.95±0.38 -0.86±1.56 17.9±0.8 Ml 16.5.95 200 3.6 4.71±0.04 1.594±0.010 3.97±0.05 l.34±0.08 1.72±0.38 1.91±1.55 17.5±0.7 Ml 16.5.95 600 3.4 4.74±0.04 1.848±0.017 3.92±0.05 2.59±0.ll 1.98±0.39 0.72±1.55 16.5±0.s* Ml 16.5.95 1020 3.25 4.94±0.04 2.232±0.025 3.79±0.04 4.85±0.15 3.92±0.41 2.42±1.57 15.4±0.6 Ml 16.5.95 1420 3.2 4.97±0.04 2.451±0.033 3.92±0.05 6.01±0.19 4.25±0.42 9.87±1.63+ 14.6±0.8 Ml 16.5.95 1520 3.15 5.04±0.04 2.434±0.034 3.72±0.04 6.10±0.20 4.94±0.43 l.57±1.61 14.0±0.7 Ml 16.5.95 1624 3.1 4.82±0.04 2.281±0.036 3.84±0.05 4.81±0.20 2.74±0.42 -0.88±1.64 14.0±1.0 NI 12.7.92 100 3.5 4.73±0.03 1.498±0.009 3.95±0.04 0.92±0.06 l.94±0.29 l.92±1.47 20.0±0.5 NI 12.7.92 240 3.5 4.78±0.03 1.517.±0.009 3.91±0.04 1.06±0.06 2.40±0.29 2.01±1.82 20.5±0.6 NI 12.7.92 400 3.5 4.79±0.03 1.66\l:l:0.012 3.83±0.04 l.83±0.08 2.54±0.30 -1.55±1.48 20.4±0.4. NI 12.7.92 600 3.45 5.12±0.03 1.865±0.016 3.63±0.04 3.39±0.10 5.84±0.32 1.11±1.53 19.3±0.5 NI 12.7.92 800 3.4 5.33±0.03 l.953±0.019 3.48±0.03 4.23±0.12 7.87±0.33 0.57±1.36 19.0±0.5 NI 12.7.92 880 3.4 5.37±0.03 l.834±0.020 3.48±0.03 3.67±0.12 8.26±0.34 1.74±1.46 18.8±0.5 NI 19.6.93 20 2.9 4.69±0.02 1.421±0.008 0.48±0.05 1.34±0.18 19.7±0.6 NI 19.6.93 200 3.5 4.68±0.02 1.571±0.008 3.95±0.03 1.19±0.05 1.41±0.18 0.07±1.20 18.2±0.5 NI 19.6.93 400 3.5 4.82±0.02 1.742±0.012 3.88±0.03 2.22±0.07 2.77±0.19 2.01±1.27 19.5±0.5 NI 19.6.93 650 3.45 5.04±0.02 1.933±0.017 3.67±0.03 3.58±0.09 5.04±0.20 ·0.04±1.17 18.7±0.6 NI 19.6.93 800 3.4 5.07±0.02 1.922±0.019 3.63±0.03 3.58±0.10 5.34±0.20 -0.87±1.19 19.4±0.5 NI 19.6.93 900 3.3 5.08±0.02 1.789±0.021 3.66±0.03 2.91±0.11 5.34±0.21 0.90±1.26 17.5±0.6 NI 24.5.95 20 0.5 4.79±0.04 1.403±0.007 3.94±0.05 0.45±0.06 1.74±0.39 -2.13±1.54 17.0±1.l NI 24.5.95 200 3.6 4.57±0.04 1.540±0.009 4.00±0.05 0.87±0.07 0.31±0.37 -2.22±1.55 16.8±0.7 NI 24.5.95 400 3.5 4.84±0.04 1.653±0.012 3.80±0.05 1.83±0.09 3.00±0.39 ·0.85±1.64 15.6±1.0 NI 24.5.95 600 3.5 4.98±0.04 1.793±0.015 3.74±0.04 2.75±0.11 4.39±0.41 1.27±1.54 16.0±0.6 NI 24.5.95 700 3.45 5.04±0.04 1.831±0.017 3.65±0.04 3.06±0.l l 5.01±0.41 ·0.90±1.55 18.1±1.0 NI 24.5.95 800 3.45 5.03±0.04 1.858±0.020 3.72±0.04 3.18±0.12 4.92±0.41 2.04±1.61 16.2±0.7 NI 24.5.95 890 3.4 4.98±0.04 l.791±0.022 3.71±0.04 2.75±0.13 4.42±0.41 -0.03±1.55 17.3±0.7 105 With respect to the isotopic composition of the light noble gases, the water masses in the Central and Southern Basins are similar. 3He/4He ratios are rather large, indicating that the water is enriched in tritiogenic 3He. However, the ratios do not fall on the tritiogenic line. Instead of this, the 20Ne/4He ratio tends to decrease with an increase in the helium isotope ratio, indicating that a terrigenic helium component is present. Because Lake Baikal is situated in a large continental rift, a crustal helium component which has a typical 3He/4He ratio of about 2·1Q-8 (Mamyrin and Tolstikhin, 1984) is more likely to be present than a mantle helium component with a typical 3He/4He ratio on the order of lQ-5. This assumption is supported by the fact that helium isotope ratios measured in hydrothermal waters from the Baikal area are affected by the presence of crustal helium (Kipfer et al., 1996). In the Northern Basin, the terrigenic component is more pronounced than in the other two basins, and the 20Ne/4He ratios are therefore smaller. 3He/4He ratios observed in the Northern Basin are also smaller than those in the other two basins. Q; Figure 6.3: Isotopic ratios of 6 0 :i: 2.510- '@go helium and neon concent- 0 ...!!. A rations measured at A A A ""'°o :i:" & positions SJ (A), CI (0) 0 0 A o" QI A ·a and NI (Ell) between I992 ::i:: 210- 6 ..... A "0 and 1995. Deviations from QI "" " " " A 0 ·-= the isotopic ratios of air- ,p:: .. , " 0 A ".E " "" 0 A saturated water (ASW) are " A caused by the decay of 6 "" 1_510- "~ 0 line) or : oA tritium (tritiogenic c111stal He by the input of radiogenic helium from the sediment 3.4 3.5 3.6 3.7 3.8 3.9 4 4.1 (crustal line). 20N e/4He The difference between the individual basins can also be seen in the 4He con- centrations. Terrigenic 4He concentrations, calculated as 4He 1er = 4Hem - 4Heeq - 4Heex at in situ water temperature are plotted in Fig. 6.4a. The distribution of 4Herer in the Southern and Central Basins is rather similar: concentrations are close to saturation at the lake surface and increase slightly with depth. The terrigenic 4He component 4Herer is largest in the Northern Basin. It reaches about 18% of the atmospheric equilibrium concentration at the bottom of the basin in 1992 (4Heeq = 4.54· l ()-8 cm3 STP g-1 for T = 3.5 °C, S = 100 mg kg-I and p = 958.9 mbar (Weiss, 1971)). 106 The 3He excess remaining after correcting for excess air, 3He;x, is shown in Fig. 6.4b. It combines the tritiogenic and terrigenic 3He components. The horizontal gradients of 3He:X between the individual basins are surprisingly small. 3He;x concentrations increase steadily from nearly zero at the surface to reach maximum values at about I 00 to 300 m above the bottom. The largest 3He;x concentrations are found in the Central Basin, where they reach 100% of 3Heeq· In all three basins, 3He;x decreases significantly in the bottom 100 to 300m. From 1992 to 1995, tritium concentrations (Fig. 6.4c) in the top 600 to 1000 m decreased by up to 4 TU. Note that radioactive decay during this 3-yr period accounts for a decrease of about 16%, or about 3 TU of the initial 20 TU. The relative shape of the tritium profile did not change significantly during this period in the Southern and Central Basins: concentrations are fairly homogeneous in the top 500 m and decrease slightly towards the bottom. In the Northern Basin, however, the decrease in tritium concentrations observed in the deep waters between 1992 and 1995 was only about l TU, which is less than expected from radioactive decay. Consequently, the vertical gradient of the 3H concentrations changed sign. Whereas in 1992 the concentration decreased slightly with depth below 300 m, the Or.11::~---r-,..-.-..,-...-,....., 200 400 ] 600 = 800 go 1000 Q 1200 1400 c) 1600 i.;;....i-.JJ:ul.-'-__...... _.,__,,_.. 0 2 4 6 8 0123456 14 16 18 20 9 3 1 14 3 1 4He [ 10" c m STP g" ] 3He' ex I 10" c m STP g" ) [TU] ter 3H Figure 6.4: Profiles of (a) radiogenic 4 He, (b) excess 3/{e ( 3He:XJ. and (c) tritium, measured at the deepest points of the Southern (A &A), Central (e &0) and Northern (11 &DJ Basins in 1992 (black symbols) and 1995 (white symbols). 3He;x is the remaining 3/{e excess after correcting for excess air at the in situ water temperature. ft combines the tritiogenic and radiogenic 3/{e components. 311e;x measured in May 1995 at position K2 (®) (at the bottom of Kukui Canyon) is included in (c ). The low value at K2 reflects the flow of surface water along the canyon. 107 0 200 400 ] 600 800 o£ip,. 0 1000 1200 1400 a) b) c) 1600 1.6 2 2.4 1.6 2 2.4 1.6 2 2.4 3H efHe 110- 61 3H e fHe 110- 61 3HefHe 110- 61 Figure 6.5: Helium isotope ratios measured at stations SJ, Cl, and NJ in 1992 (A), 1993 (e), and 1995 (D). 1995 observations show an increase of about 2 TU below 600 m. Since in 1992 the 3H concentrations in the surface layer were larger than those in the deep waters, the reversal of the gradient observed between 1992 and 1995 implies that significant quantities of surface water from either the Central or the Northern Basin must have been transported into the depths of the Northern Basin. Although the temporal variation of the helium isotope concentration is fairly complex, the vertical distribution of the 3He/4He ratios remained remarkably constant in all basins between 1992 and 1995 (Fig. 6.5). At the water surface, the 3He/4He ratios are close to the solubility equilibrium (Req = 1.360· lQ-6 (Benson and Krause, 1980)) and increase steadily with depth. Maximum values are reached at the bottom in the Southern Basin and at 100 - 300 m above the bottom in the Central and Northern Basins. 6.3.3 Helium Isotopes in Hydrothermal Springs In order to determine the isotope ratio of the injected terrigenic helium component, water from hydrothermal springs near the shore of the Northern Basin was analysed for helium and neon. A first estimate of the 3Hef4He ratio of the injected terrigenic component was calculated by Kipfer et al. (1996) based on two samples taken at the Khakusy hot spring (see Fig. 6.1). Additional data are given in Table 6.2. 108 In all hot spring samples, the 4He concentration is very large. It varies by a factor of 50 between Kotelnikovsky (highest value) and Zmeinnyi Istochnik (lowest value). The absence of a comparable neon excess implies that the helium is not of atmospheric origin, but represents the terrigenic helium component. In contrast, the isotope ratios vary by only a factor of 10. The lowest 3ffe/4He ratios were measured in samples with the highest 4He concentrations (Kotelnikovsky), and the highest 3ffe/4He ratios were measured in samples with the lowest 4He concentrations (Zmeinnyi Istochnik). All 3Hef4He ratios are significantly lower than the atmospheric ratio, but higher than that of helium from the stable continental crust (Rcr"' 2· lo-8; (Mamyrin and Tolstikhin, 1984)). This indicates that, in addition to the helium from the Earth's crust, small amounts of mantle helium may also be present. The measured helium isotope ratios are in excellent agreement with the data published by Polajak et al. (1992). Unfortunately, no information on the discharge of the different hydrothermal springs is available, so that the average composition of the helium injected into Lake Baikal cannot be calculated by weighting the measured ratios with the flow rates. The mean value of the measured ratios (R1er"' 2.2 ± 1.8· I0-7; the samples from Khakusy were treated as a single value) was therefore used in the following discussion. A similar value was reported by Craig and Lupton (1978) for the helium injected into Lake Tanganyika ('" 4·1Q-7). Tabk 6.2: Helium and neon concentrations measured in hydrothermal springs in the vicinity of the Northern Basin of Lake Baikal. 4He concentrations and 3Hef4He ratios were corrected for air contamination using 20Ne excess0 at the in situ temperature. Name Position Date T[OC] 4He 3Hef4He 20Ne [J0-6cm-3sTP ~-11 [ 10-7 i [I0"7cm·3sTP g-11 Khakusyb Ksy 9.7.92 46 7.89 ± o.os 1.78 ± O.Ql 1.68 ± 0,02 Ksy 9.7.92 46 7.98 ± 0.06 1.79 ± O.ot 1.70 ± 0.02 FrolikhaBay FB 12.7.93 36 23.63 ± 0.09 1.91 ± 0.02 S.45 ± 0.04 Zmeinnyi Istochnik z 10.11.94 so 2.88 ± 0.02 4.68 ± 0.02 l.35 ± O.ot Kotelnikovski:: Kot 27.5.95 80.5 103.65 ± 0.84 0.50± 0.01 3.28 ± O.Q3 a atmospheric equilibrium concentrations of neon is between l.42·10--7 cm-3 STP g-1 (al T= 36 °C) and I.06·10--7 cm·3 STP g-1 (at T= 80.5 °C) (Weiss, 1971). b values from Kipfer et al. (1996). 109 6.4 Discussion 6.4.1 3ff.3ffe Ages The 3ff.3He ages were calculated according to Eq. (6.3). The 3ff.3He age yields merely an estimate of the true age, defined as the time since the water was last in contact with the atmosphere (Mamyrin and Tolstikhin, 1984; Torgersen et al., 1977). Since the 3ff.3He age is a non-linear function of 3H and 3He1r1, mixing may result in a descrpancy between it and the true age, especially if 3H-gradients are large and if tritium concentrations in the lake are not at steady state (Jenkins and Clarke, 1976; Aeschbach-Hertig, 1994). In particular, the non- linearity of Eq. (6.3) gives rise to two specific effects. (1) The 3H-3He age of a mixture of two water masses with different 3H concentrations is biased towards the 3H.3He age of the component with the higher 3H concentration. (2) Because of the logarithm in Eq. (6.3), the 3H-3He age always overestimates the true age of a mixture of two water masses. In the case of Lake Baikal, these effects cancel each other out to some extent. Except for the 3H profile measured in the Northern Basin in 1995 (Fig. 6.4c), 3ff concentrations decrease with increasing depth, so that the former effect leads to a slight underestimation of the true age of the deep waters. Furthermore, 3ff gradients in the lake are small and deviations due to the non-linearity of Eq. (6.3) are therefore also expected to be small. The 3H-3He age and the true age of a mixture of Baikal surface water (3H =18 TU; 3H-3He age 't = 0) and deep water (3H = 16 TU; 't = 10) are compared in Fig. 6.6. The figure illustrates that, in the case of Lake Baikal, the 3H-3He age generally yields an overestimate of the true age of the mixed water mass. However, since the discrepancy does not exceed 10%, the 3H- 3He age can be used as a first-order estimate to quantify mixing processes in the lake. 10 . 8 .. :g 6 4 Figure 6.6: Comparison between the 3H.3He age (dotted ~ .. line) and the true age (solid line) of a mixture of Baikal 2 surface water (3H = 18 TU; 3H-3He age i- = 0 yr) and deep water (3H = 16 TU; i- = JO yr). f1 is the relative 0 proportion of surface water. 0 0.2 0.4 0.6 0.8 1 Tl 110 3Heiri was calculated from Eq. (6.4) using the 3He/4He ratio of the injected terrigenic 7 component determined above (R1,, = 2.2· 10- ). The error in the 3ff-3He age due to the uncertainty in R1er is negligible. A maximum terrigenic 4He excess of about 18% was observed at the bottom of the Northern Basin in 1992. The corresponding terrigenic 3ffe excess is approximately 2.7%, compared to the tritiogenic 3He excess of about 58%. With R1er= 4· 10-7, the terrigenic 3ffe excess would be 5.3%, the tritiogenic component would therefore be 55.4%, and the 3H-3He age would decrease from 10 to 9.7 yr. 3ff-3He ages in the three basins from 1992 to 1995 are shown in Fig. 6.7. Again, the similarity of the curves in all three basins is remarkable: the 3H-3He ages increase steadily with depth from values close to zero at the lake surface to their maxima at about 100 300 m above the lake bottom. The greatest 3H-3He ages found were 13 - 17 yr in the Southern Basin, 16 - 18 yr in the Central Basin and IO 11 yr in the Northern Basin. The maximum value found in the Central Basin agrees well with the age of 18 yr reported by Craig (1989). In the bottom waters of all basins, the increase in 3ff-3He age with depth is halted (Southern and Central Basins) or even reversed (Northern Basin, sometimes in the Central Basin). The particular shape of the 3H-3He age curve and its temporal evolution implies (1) that the bottom layer is renewed more efficiently than the water masses found directly above the bottom layer, and (2) that the intensity of this renewal differs from basin to basin and from year to year. o·.k'I,...,.,..,..,...,...,.,..,..,....,.,.,.,..,....,.,., 200 400 ] 600 ..s 800 g 1000 1200 1400 1600~~~~~~~~~ 5 10 15 5 10 15 5 10 15 3H-3He age [yr] 3H-3He age [yr] 3H-3He age [yr] Figure 6.7: 3ff.3He ages calculated from Eq. (6.3) at stations SJ, Cl, and NI in 1992 (ll.), 1993 (e). and 1995 (DJ. ll1 In the Southern Basin, the 3H-3He age remains fairly constant in the water column above 1000 m depth during the period 1992 - 95. In contrast, the 3H-3He age in the bottom layer (the lowermost 200 m) increases steadily from 13 yr in 1992 to about 17 yr in 1995, i.e. by roughly one year per year, implying practically no renewal of the bottom layer during this time period. In the Central Basin, variations of 3H-3He age with time and depth are more complex. Minor variations in the upper 600 m may be caused by convection triggered by the thermal bar that develops along the shore in early summer (Shimaraev et al., 1993). However, the vertical resolution of the data is not sufficient to identify specific mixing processes. In the bottom layer, a distinct decrease in 3H-3He age with depth was observed in 1992 and 1995, but in 1993, the 3ff-3He age was approximately constant below 1300 m. It appears that, on average, mixing events keep the 3ff-3He age at near steady-state values. In the Northern Basin, temporal changes in 3ff-3He age occur mainly below 600 m depth. Near the bottom, the 3H-3He age decreased from 10 yr in 1992 to 8.5 yr in 1993, but remained fairly constant from 1993 to 1995. In summary, between 1992 and 1995, the 3ff-3He age in the surface and intermediate layers of all basins remained approximately constant, i.e., the increase in 3ffe concentration due to the radioactive decay of tritium was balanced out by the decrease due to the flux of excess 3He to the atmosphere by vertical mixing and gas exchange. A different picture was found in the bottom layers. In the Southern Basin, the 3ff-3He age increased steadily during the observation period, indicating that the renewal of the bottom layer was proceeding less efficiently than in previous years. In the Central and Northern Basins, renewal of the bottom layer varied from year to year, but on average the 3ff.3He age was practically constant. 6.4.2 Comparison with CFC-12 Ages The relative vertical distribution of 3H.3He ages observed between 1992 and 1995 is similar to the distribution of the CFC-12 ages in 1989 reported by Weiss et al. (1991), although the 3H-3He method yielded larger values than the CFC-12 method. Maximum CFC-12 ages do not exceed 14 yr in the Southern Basin, 16 yr in the Central Basin and 9 yr in the Northern Basin. Rather than indicating a real age increase from 1989 to 1992, the differences between 3H.3He and CFC-12 ages are most likely caused by methodological factors. For instance, it was shown in section 6.4.l that due to the non-linearity of Eq. (6.3), the mixing of two water masses may result in the 3ff.3He age overestimating the true age by up to about 10%. On the other hand, the CFC-12 ages given by Weiss et al. (1991) were derived from CFC-12 112 concentrations which had been corrected for an undersaturation measured in the upper 250 m layer. This correction might result in a discrepancy between CFC-12 age and true age. Furthermore, as in the case of the 3H-3He age, mixing may cause the CFC-12 age to deviate from the true age due to the non-linear evolution of the atmospheric CFC-12 concentration. The observed discrepancy between the CFC-12 and 3H-3He ages is, to a large extent, likely to be due to a combination of these various effects. 6.4.3 Deep-Water Renewal Rates A detailed evaluation of the development of 3H-3He ages in terms of deep-water renewal rates requires the application of a dynamic mixing model of the lake and information on the tritium input from the atmosphere during the last 30 years (Peeters et al., 1997). However, even a simple approach based on volume-weighted mean 3H-3He ages below 250 m depth provides a reasonable estimate of the mean residence time of the waters in this layer (Table 6.3). Deep- water renewal rates are about 10% yrl in the Southern and Central Basins, and about 15 % yrl in the Northern Basin. Corresponding mean vertical exchange velocities across the 250 m isobath are 71 81 m yrl in the Southern Basin, 73 77 m yrl in the Central Basin and about 58 - 61 m yrl in the Northern Basin. These values are in good agreement with the mean value of 75 m yrl given by Weiss et al. (1991). In comparison, the vertical velocities of 30 to 150 m yrl at 400 m depth reported by Killworth et al. (1996) seem rather large. Table 6.3: Mean 1fl.1He age below 250 m depth and mean vertical velocity vz of water across the 250 m isobath. The calculation of the volume- weighted mean water ages is based on samples from stations SI, S2, Cl, C2, C3, C4, NI and N3 (see Fig. 6.1). The water volumes below 250 m depth are taken from Shimaraev (1994). ii is the mean depth of the basin below 250 m. Basin 'ii [m] year age[yr] Vz[m yrl] 1992 9.64 ± 0.47 SL! South 781 1993 9.86 ± 0.82 79.2 1995 I I.OJ ± 0.51 70.9 1992 10.54 ± 0.38 76.5 Central 807 1993 11.10 ± 0.52 72.7 1995 10.47 ± 0.55 77.0 1992 7.19 ± 0.85 61.0 North 438 1993 7.62 ± 0.73 57.6 1995 7.52 ± 0.80 58.3 113 Note that in the Southern Basin, the mean 3H.3He age of the deep water increased steadily from 9.6 yr in 1992 to 11 yr in 1995, which caused the vertical exchange velocity to drop from 81 m yrI to 71 m yrl. This is consistent with the observation made above that the renewal of the deepest 200 m of the basin was slight during the observation period. 6.4.4 Oxygen Consumption Rate The 3H.3He age provides a natural time scale for the determination of oxygen consumption rates in aquatic systems (e.g. Jenkins, 1976, 1987; Aeschbach-Hertig et al., 1996a). Figure 6.8 shows the oxygen saturation anomaly calculated at the in situ temperature plotted against 3H.3He age below 200 min all three basins for the years 1993 and 1995 (the oxygen data from the 1992 cruise are too sparse to be employed here). Since the regression shows neither significant differences between the basins nor from year to year, the average oxygen consumption rate per unit volume, given by the slope of the regression line, seems to be constant in space and time: J02 =(4.5 ± 0.3) µmol J-l yrl. To express the consumption rate per unit lake area, l0z has to be multiplied by the basin-specific mean depths. Due to its reduced mean depth, the oxygen consumption per unit area would be smaller in the Northern Basin than in the other basins. Oxygen is consumed both in the water column and at the sediment surface. Thus, the rate calculated above combines both the volume ( 102 ,v) and the areal ( l0z,A) oxygen sinks (Livingstone and Imboden, 1996). A rough estimate of the volume oxygen sink, J02,v = (4.2 ± 0.5) µmoJ 1-l yrl, is yielded by neglecting the data from the 200 m thick bottom layer of each basin where the ratio of sediment area to water volume is particularly large. -20 Figure 6.8: Correlation .. 0 -30 between oxygen deficit and ' ' 3H.3He age in the Southern ,,...... -40 '~) .. (•), Central (0) and ...... ' Northern (EB) Basins in 0 ..... , -so 1993 and 1995. The slope ], 0 -60 0 of the regression line de- ON 'Aj fines the oxygen consump- Both rates are similar to the values determined by Weiss et al. (1991) (4.5 µmol J-1 yrl below 250 rn; 3.8 µmol 1-1 yr' at mid depth below 250 m) and by Killworth et al. (1996) (4.0- 5.8µmol1-1 yr'). Funhermore, both values coincide remarkably well within the error associated with the linear regression, indicating that oxygen consumption in Lake Baikal is dominated by the volume sink rather than the areal sink. 6.4.5 4ffe Flux from the Lake Bottom A similar approach can be adopted to determine the flux of helium from the lake bottom. There is an excess of 4He in the deep water of all three basins (Fig. 6.4a). Excess 4He plotted against the 3H-3He age is shown in Fig. 6.9. As a first approximation, the 3H-3He age is assumed to be independent of the flux of terrigenic helium from the lake bottom (R1,, = 2.2· I0-7), and thus the 4He-flux per unit volume from the sediment (J He) can be estimated as the slope of the regression curve. Multiplying by the mean basin depth h yields the He flux per unit area, FHe (Table 6.4). Helium fluxes in the Southern and Central Basins are similar (FHe"' 1.3· lQll atoms m-2 s-1), while in the tectonically most active Northern Basin the flux is significantly larger Table 6.4: Mean volumetric oxygen consumption rate lo2 ':_nd mean helium flux from the sediment FHe· h is the mean depth of the basin below 200m. Basin h 102 FHe [m] [l!mol 1-I :z:r-l] [atoms m-2 s-1] South 844 (l.38 ± O.l l)·JOl l Central 854 4.5 ± 0.3 (L24 ± 0.14)·1011 North 576 (2.79±o.1spo•1 115 10 ...... -.- ...... - ...... - ...... ,...... -.-.,...... , ...... ,....., Figure 6.9: Correlation between 4He excess and 1H- 1He age in the Southern (.IJ,.), 8 IB Central (0) and Northern (81) Basins for the years 1992, 6 1993, and 1995. The 0 0 regression lines have been 4 forced through the origin. The slopes of the curves define the volumetric helium 2 accumulation rates. In the Southern and Central Basins O'-"'::.;.,.....i.;;;_._1...... 1_.__.__._...... _...... _.__,_...... ,'--'-_._....._. these are quite similar 0 5 10 15 20 (I.92·JO-IO and l.7J.J0-10 3 3 H- He age [yr) cmlSTP g-1 yrl, respectively). Jn the Northern Basin the rate is higher (5.69· JO"-IO cmlSTP 8-1 yrl). 3 Multiplying the helium flux by the terrigenic isotope ratio R1er yields the flux of He. Comparing this with the average geothermal heat flow in Lake Baikal (71 ± 21 mW m-2 (Golubev et al., 1993)), we obtain a heat/3He ratio of I. I to 2.6· 10-6 J atom-I. In submarine hydrothermal systems, heat/3He ratios range from 0.04 to 2.6-1 Q-6 I atom-I (Lupton et al., 1989). Hydrothermal activity combined with high advective geothermal heat flow is known to occur in the Northern Basin (Golubev et al., 1993). Kipfer et al. (1996) traced the helium- rich plume of a sub-aquatic spring at Frolikha Bay that was flowing towards the deep part of the basin. Since helium concentrations measured in the various springs in the vicinity of the Northern Basin vary by a factor of 50 (Table 6.2), it is not possible to calculate the water discharge of the springs and the total input of helium into the basin. 6.4.6 A Simple Model for Advective Bottom-Water Ventilation The peculiar structure of the vertical temperature, salinity, helium and tritium profiles (Figs. 6.2 and 6.4) indicate that the renewal of the bottom water must be mainly the result of large- scale advection. This has been described in detail by Hohmann et al. (1997a) and will be summarised only briefly here. River-induced deep-water exchange in the Central Basin is associated mainly with the Selenga River, the major inflow entering the lake at the boundary between the Southern and 116 Central Basins. Kukui Canyon cuts into the northern slope of the Selenga Delta, leading from near the shore to the deep part of the Central Basin. During April and early May, cold and relatively saline water from the Selenga River mixes with lake surface water near the river mouth, forming a relatively dense plume that flows along and down the canyon to the deepest part of the lake (Hohmann et al., 1997a). The 3ffe concentration in a water sample taken in the canyon (position K2, Fig. 6.1) on May 15, 1995 confirms the flow of 'young' surface water along the bottom of the canyon (Fig. 6.4b). The 3He concentration at K2 is still close to the atmospheric equilibrium value and the 3He excess is significantly smaller than that at a similar depth at Cl. An analogous phenomenon is to be expected for that part of the Selenga inflow entering the Southern Basin. In fact, distinct temperature and salinity signals recorded in June 1993 and May 1995 (Hohmann et al., 1997a) indicate that water from the Selenga River - or another inflow - reaches the very bottom of the basin. However, the signals are restricted to the small 2 - 5 m bottom layer at the deepest part of the basin. So far, in contrast to the Central Basin, we have not been able to trace the penetration of a plume of river water into the Southern Basin directly. Possibly, the absence of a topographic structure comparable to the Kukui Canyon renders detection more difficult. In the Northern Basin we found a distinct cold bottom boundary layer in May 1995 with a volume of 40 - 80 km3 covering an area of approximately 4000 km2. A low- temperature signal was also recorded in June 1993, but at that time it was restricted to the very deepest part of the basin. From a detailed set of CTD profiles (Hohmann et al., 1997a) we were able to identify Academician Ridge - separating the cold, relatively saline water of the Central Basin from the warmer and slightly less saline water of the Northern Basin - as the origin ofthe cold bottom water detected in 1995. At the ridge, horizontal mixing produces water that can sink on either side of the sill. While in the Central Basin the water mass remains at an intermediate depth, in the Northern Basin it sinks to the bottom. The following simple model is based on these phenomena. Its application to 3He serves to estimate the total volume of water flowing to the bottom by large-scale advection. The mass balance of 3He in the bottom layer of mean thickness ii is described by the equation A'(3J..J,. \ (3 3 ) K a(3He) 3 = kad Head - He R ·Jn ~ H (6.5) :i..:.:::ldt v v + 1er eh+ ·-- dZ + ')..,. with mean concentrations in the bottom layer [cm3 STP g-1) He concentrations in the advective plume [cm3 STP g-1] advective exchange coefficient [yrl] 117 helium flux from the lake bottom per unit volume [cm3 STP g-1 yrl] terrigenic helium isotope ratio [ - ] coefficient of vertical diffusion [m2 yrl] mean thickness of bottom layer, i.e. volume divided by isobath area at upper boundary of layer [m] z vertical coordinate, positive upwards [m] A.= 0.05576 yrl tritium decay constant (Unterweger et al., 1980). The terms on the right-hand side describe the effects of the following processes on the 3He balance: input by advection; flux of terrigenic 3He from the lake bottom; turbulent diffusion through the upper boundary; and production by the radioactive decay of tritium. 3 Solving Eq. (6.5) for kadv and replacing d(3He)/dt by t. He/t.t yields 3 K a(3He) 3 t. He ~·--+R ·JH H--- _ h dz /ere +A.· ,1t (6.6) kadv - (3 He-3Headv ) Equation (6.6) is evaluated separately for each basin. The bottom layers are defined as follows: Southern Basin, depth> 1200 m; Central Basin, depth> 1400 m; Northern Basin, depth > 700 m; i.e. each bottom-water reservoir is approximately 200 m thick. The following approximations are made: The 3He concentration in the bottom layer is calculated as the average of the volume- weighted means from two consecutive expeditions (i.e., 199211993 and 1993/1995). The vertical gradient of 3He is approximated by linear regression of the data from the deepest part of each basin. Note that 3He concentrations are approximately constant in time. The helium concentration in the advective plume is set equal to the saturation value at 3.5 °C (3Headv"' 3Heeq(3.5 °C) =6.17·10-14 cm3 STP g-1 (Weiss, 1971; Benson and Krause, 1980)), implying that the plume originates from a well-mixed surface layer which is in atmospheric equilibrium. Because the effect of entrainment in the water column on 3Headv is neglected, the absolute value of the denominator of Eq. (6.6) tends to be too large. Thus, the resulting kadv value represents a lower limit to the overall advective transport into the bottom layer. Since information on vertical diffusion is lacking, we use Kz = lfr4 m2 s-1 in all three basins. This is a typical value for the weakly-stratified hypolimnia of deep lakes (Wllest, 1987). 118 The different terms ofEq. (6.6) and the resulting exchange rates are listed in Table 6.5. The absolute advective transport, fladv, is calculated as the product of kadv and the volume of the respective bottom layer. Table 6.5 shows that the mass balance of3He is dominated by the decay of tritium; its contribution exceeds the effect of diffusive mixing by a factor of two. Changing Kz by ± 50% causes kadv and Qadv to change by 20% or less. The terrigenic 3He flux is negligible. In the Southern Basin the advective flow is virtually zero, and between 1992 and 1993 even negative. However, in 1992 there was only one single sample collected below 1200 m, and the volume-weighted mean 3He concentration is therefore poorly defined. Nevertheless, the calculations confirm the previous observation that bottom-water renewal was slight in this basin between 1992 and 1995. In the Central Basin, a total surface water volume of 10 to 20 km3 yrl is required to explain the observed mean 3He concentrations in the bottom layer. Hohmann et al. (1997a) have shown that a total volume of about 2.5 km3 yrl from the Selenga River has the potential to reach the bottom of the Central Basin. An additional 7.5 to 17.5 km3 yrl of water from the surface mixed layer are therefore necessary to produce the required fladv· According to a salinity balance, the descending plume observed in Kukui Canyon at 400 m depth (position K2) consists of approximately 1/3 Selenga water and 2/3 lake surface water (Hohmann et al., 1997a). Thus, it is likely that bottom-water renewal in the Central Basin is to a large extent accomplished by a mixture of river water and lake surface water from the delta region. The largest advective transport about 150 km3 yr1 between 1992 and 1993 and 80 km3 yr! between 1993 and 1995 - is found in the Northern Basin. The latter value coincides remarkably well with the volume of the cold bottom layer detected in tbe Northern Basin in spring 1995 (Hohmann et al., l 997a). In contrast to the Central and Southern Basins, bottom-water renewal in the Northern Basin is primarily induced by the mixing of water masses from the Central and Northern Basins at Academician Ridge. The fluxes into the 200 m thick bottom layers of each individual basin derived from the vertical velocities given by Killworth et al. (1996) are on the same order of magnitude. Independent confirmation of the results obtained from Eq. (6) by calculating the mass balance of other tracers such as CFCs would be helpful. Due to the noisiness of the data set, the uncertainties in the spatial and temporal gradients of 4He and tritium are too big to allow their mass balances in the bottom layers to be evaluated meaningfully. Table 6.5: Advective bottom-water renewal rates kadv calculated from Eq. (6.6). The bottom layers are defined as_jollws: Southern Basin, depth> 1200 m; Central Basin, depth > 1400 m; Northern Basin, depth > 700 m. Vis the volume and h the mean depth of the 200-m bottom layer. Kz is the coefficient of vertical diffusion and Qadv the absolute advective flow. Jse is the 4He flux from the lake bottom and R1er the terrigenic helium isotope ratio. basin v h year !J.t 3He 3H &. 6.5 Summary and Conclusion A detailed analysis of He, Ne and 3H in Lake Baikal between 1992 and 1995 shows that the water masses from the Central and Southern Basins have a similar composition and vertical structure. In the Northern Basin, an injected terrigenic helium component is resulting in high 4He concentrations and relatively low 20Ne/4He ratios. The 3He/4He ratio of the injected terrigenic component is remarkably similar to the value reported by Craig and Lupton ( 1978) for the helium injected into Lake Tanganyika. The geochemical implications of this observation remain unclear. Deep-water renewal rates calculated from volume-weighted means of 3H.3He ages below 250 m depth are approximately 10% yr-l in the Southern and Central Basins and about 15% yr-1 in the Northern Basin. In the Southern Basin, the mean 3H-3He age below 250 m depth increased steadily from 9.6 yr in 1992 to 11 yr in 1995. Apparently, during these years vertical exchange was not efficient enough to prevent the accumulation of tritiogenic 3He. Over longer periods of time, the mean 3H.3He age in the Central and Northern Basins seems to be fairly constant. The distinct decrease in 3He concentration and in 3H-3He age with depth in the lowermost 100 to 300 m implies that the renewal of the bottom layer is the result of large- scale advection. This is not necessarily the case in the water column above the bottom layer, where the distribution of these tracers could be explained by the effect of vertical diffusion and of the radioactive decay of tritium. In the Southern Basin the advective transport of surface water to the bottom layer, calculated from the 3He balance for the period 1992 to 1995, is practically zero. However, the concentrations measured in 1992 indicate that at some point previously the renewal of the bottom layer must have been more intense. In the Central Basin, advective bottom-water renewal occurs at rate of between 10 and 20 km3 yrl, and in the Northern Basin between 80 and 150 km3 yrl. In the Central Basin, and most likely in the Southern Basin, bottom-water renewal is to a large extent controlled by the discharge rate and the water temperature of the Selenga (Hohmann et al., 1997a). Deep-water mixing in the Northern Basin is controlled mainly by the magnitude of lateral mixing at Academician Ridge between water from the Central and Northern Basins (Hohmann et al., l 997a). The mixing of water masses at the sill provides a significant supply of cold and 'young' surface water flowing to the bottom of the Northern Basin, where it forms a cold bottom boundary layer as observed in spring 1995. The large advective bottom-water renewal rates in the Northern Basin are also reflected in the distribution of tritium (Fig. 6.4c): between 1992 and 1995, the decrease in 3H concentration below 600 m was much smaller than would be expected from radioactive decay alone. This can only be explained by the downward mixing of large quantities of surface water to the deep part of the basin. 121 In the Southern Basin, the intensity of deep-water renewal now seems to be smaller than the Jong-term average. At the present time we are not able to decide whether the steady increase in mean 3H-3He age below 250 m is the manifestation of a Jong-term saw-tooth fluctuation as observed in some Swiss lakes (Livingstone, 1993, 1997) or whether it indicates that the basin is shifting towards a permanent state of reduced deep-water renewal. The latter could have severe implications for the lake's unique ecosystem. Acknowledgements Among the many people who helped with the planning, the preparation and the realisation of our research on Lake Baikal, we would like to express our special thanks to our colleagues from the Limnological Institute in Irkutsk, especially to its director M. Grachev, but also to N. Cherepanova, N. Granin, A. Zhdanov and T. Khodzher and the crew of the RV Vereshchagin. Michael Schurter was responsible for the technical equipment on all the cruises. Laboratory work would not have been possible without the support of Rainer Wieler, Ors Menet, Stefan Thiirig, and Werner Aeschbach-Hertig. David Livingstone helped to improve the language of the final text. Ship time and support were provided by the Baikal International Center of Ecological Research (BICER) and the Limnological Institute of the Siberian Division of the Russian Academy of Sciences. This research was made possible by financial support from the Swiss Federal Institute of Environmental Science and Technology (EAWAG), the Swiss Federal Institute of Technology (ETH) and the Swiss Federal Office for Science and Education (BBW). 7. Summary and Outlook Based on the theoretical concepts to describe the local, vertical stability of a water column and neutrally buoyant transport in cold, deep freshwater lakes (Sect. 2), and on the analysis of a detailed set of CTD profiles and chemical data (L. Sigg, unpublished data), processes of deep-water renewal could be identified and were discussed in Sects. 3 and 4. Deep-water renewal rates were determined from the results of noble gas analysis in Sect. 6. In this final section, these results will be summarised and a brief outlook on possible future research activities will be given. 7 .1 Deep-Water Renewal in Lake Baikal Deep-water ventilation in Lake Baikal occurs in spring and late autumn. It is the result of the delicate interplay between chemical and physical factors and is most likely to occur in regions where water masses with different (0,Sc)-characteristics meet horizontally. From the distribution of temperature and salinity two important sites of deep water formation were identified: (I) the Selenga Delta between the Southern and the Central Basins; and (2) the Academician Ridge, separating the Central and Northern Basins. Deep-water renewal rates calculated from volume weighted means of 3H-3He ages below 250 m depth are approximately 10 % yrl in the Southern and Central Basins and about 15 % yrl in the Northern Basin. These rates include the effects of both diffusion and large-scale advection. The peculiar structure of the vertical temperature, salinity, helium and tritium profiles indicate that the renewal of the bottom water must be mainly the result of large-scale advection. The share of advective bottom-water renewal was calculated from the 3He mass balance in the deepest 200-m layer of each individual basin. 123 Central Basin: River-induced deep-water renewal is important in the Central Basin where it is associated mainly with the Selenga River, the major inflow entering the lake at the boundary between the Southern and Central Basins. Kukui Canyon cuts into the northern slope of the Selenga Delta, leading from near the shore to the deep part of the Central Basin. In spring, the plume of the Selenga could be identified flowing through Kukui Canyon to the deep part of the basin. Important characteristics of the Selenga with respect to deep-water formation are its relatively large salinity compared to the lake water, and its large annual variation in discharge rate and in water temperature. The latter leaves two short time periods in April and early May and in late October when the river water is not yet too warm and discharge is already large enough to form a relatively dense plume that flows along and down the canyon. During these periods, the discharge of the Selenga into the Central Basin is approximately 2.5 km3. On its way along the canyon, the river water entrains significant quantities of lake water from different depths. The entrainment factor determined from the ionic salinity signal measured within the canyon increases from 2 - 3 at 400 m depth to 6 - 7 at 600 m depth. From the enhanced salinity of the river plume near Boldakovo at about 1000 m depth and at the bottom of the basin, the "final" entrainment factor was estimated at about 50. According to the 3He mass balance in the deepest 200 m layer, the advective bottom water renewal by surface water is approximately 10 - 20 km3 yr!. Therefore, the total volume of Selenga water that has the potential to reach the bottom of the Central Basin (2.5 km3 yrl) would have to grow by a factor of 4 - 8 due to the entrainment of surface water in the shallow delta region. Considering the entrainment factors derived from the ionic salinity measured in Kukui Canyon, this would appear a possible scenario. As the river water becomes warmer in late spring, it ceases to sink to the deep part of the basin and flows at the lake surface along the eastern shore where it initiates the thermal bar. At Boldakovo, creation and evolution of the thermal bar is controlled by the flow of warm Selenga water along the shore. According to chemical analysis, the water on the inshore side of the thermal bar consists of 20 - 50% Selenga water. This agrees well with the volume fraction of river water on the near-shore side of the thermal bar calculated from the ionic salinity (35-55%). Vertical exchange near the thermal bar is restricted to the uppermost 300 m. Despite its high salinity, the water near the shore is too warm to sink to the bottom of the Central Basin. 124 Southern Basin: The similar distribution of temperature and ionic salinity in the Southern and Central Basins suggests that river induced deep-water renewal is also important in the Southern Basin. In fact, since the water temperatures are slightly lower and concentrations of dissolved ions are slightly larger in the Southern Basin than in the Central Basin, the potential of Selenga water to sink to the deep part of the Southern Basin is even bigger than to sink to the deep part of the Central Basin. However, except for some distinct temperature and ionic salinity signals which were restricted to the small 2 - 5 m bottom layer at the deepest part of the basin, there was no indication of significant deep-water renewal between 1992 and 1995. So far, in contrast to the Central Basin, we have not been able to trace the penetration of a plume of river water into the Southern Basin directly. The results of noble gas analysis confinn the absence of significant deep-water renewal between 1992 and 1995. The mean 3ff-3He age in the Southern Basin below 250 m increased steadily from 9.6 yr in 1992 to 11 yr in 1995. Apparently, vertical mixing was not efficient enough to prevent the 3He concentration and the 3ff.3He age from growing during these years. According to the mass balance of 3He in the 200-m bottom layer, advective bottom water renewal was practically zero. However, the concentrations measured in 1992 indicate that at some point previously the renewal of the bottom layer must have been more intense. 20 40 60 80 Time [d] since December 1, 1995 Figure 7.1: Preliminary results from a thermistor chain moored at the deepest part of the Southern Basin since December 1995 (Figure provided by 0. Kocsis). Average temperature at 1400 m depth is about 3.4 °C. The three profiles are shifted by 0.1 °C. 125 At the present time we are not able to decide whether the steady increase in mean 3H- 3He age below 250 mis the manifestation of a long-term saw-tooth fluctuation as observed in some Swiss lakes (Livingstone, 1993, 1997) or whether it indicates that the basin is shifting towards a permanent state of reduced deep-water renewal. Preliminary results from a thermistor chain which has been moored at the deepest part of the Southern Basin since December 1995 (Fig. 7.1) indicate that a major deep-water renewing event occurred at the beginning of January 1996. At 1390 m depth, temperature decreased by approximately 0.2 °C within a period of a few hours. At 1290 m depth, a temperature decrease of 0.1 °C was observed one day later. Smaller temperature fluctuations were registered at the end of January and at the beginning of February, 1996. Another temperature decrease similar to the one registered at the beginning of January occurred in June 1996. Unfortunately, no salinity measurements are available from these periods and thus it is impossible to say whether the intruding cold water masses originate from the Selenga. Northern Basin Deep-water renewal in the Northern Basin is triggered by the small difference in temperature and ionic salinity between the Central and Northern Basins. At Academician Ridge, horizontal mixing of the cold and "saline" water of the Central Basin and the warmer and slightly less saline water of the Northern Basin results in dense water that can sink on either side of the sill. While in the Central Basin the water masses remain at intermediate depth, in the Northern Basin they sink to the bottom. From the available CTD data it appears that this deep-water forming process takes place over a wide area of the Academician Ridge. Due to the poor temporal resolution of the data, any information on the length of the deep-water formation period is missing. In spring 1995, the sinking water masses at the Academician Ridge caused a cold bottom boundary layer in the Northern Basin, covering an area of approximately 4000 km2 and holding a volume of 40- 80 km3. In June 1993, the cold bottom water was found at the very deepest part of the basin, only. According to the mass balance of 3He in the 200-m bottom layer of the Northern Basin, advective bottom water renewal is in the order of 100 km3 yrl. This value is in good agreement with the volume of the cold bottom boundary layer detected in spring 1995. 126 7 .2. Ion Budget Given the extreme sensitivity of deep-water formation with respect to the salinity gradients in the lake, it is easy to see that deep-water renewal may be strongly influenced by possible long-term changes in the salinity of Lake Baikal. Compared to the large volume of water, the annual exchange of water and salt is very small (Tab. 1.1). Therefore, spectacular changes in the physical and chemical characteristics of the lake are not to be expected. It is also very difficult to detect small residual fluxes in the total salt balance on a year-to-year basis. The mass balance of the dissolved ions in Lake Baikal is: dSc,lake = ..!..(~ nin. sin. ~Qout. sout) (7.1) dt v "-' ~ C,l k J C,J ' with Sc.lake: the mean ionic salinity of the lake water [mg kg-l]; s:~i: the mean ionic salinity of inflow i [mg kg-I]; S%,j: the mean ionic salinity of outflow j [mg kg-I]; Q;n: the mean discharge of inflow i [km3 yrl]; Q';ut: the mean discharge of outflow j [km3 yrl]; V= 23015 km3 the lake volume. Discharge and ionic salinity of the major inflows and of the only outflow, the Angara, are summarised in Tab. 7.1. Assuming that the water budget of the lake is in balance, the difference between total inflow and outflow must be accounted for by precipitation and, to a small extent, by the discharge of hydrothermal springs. Unfortunately, the data base on the discharge rates of the various inflows is rather poor. In particular, data on annual discharge and mean ionic salinity are available only for the seven major rivers entering the lake (Tab. 7.1). However, these rivers account for approximately 79.2% of the total annual river inflow ( Q;n) and we will therefore approximate the mean ionic salinity of the total inflow by the volume weighted mean of the ionic salinity of these rivers (Sc.in ""112 mg kg-'). Assuming that the ionic salinity of the Angara is equal to the mean ionic salinity of the lake water, Eq. (7.1) can be formulated as follows: dSc lake 1 ( - ) --&- =V '4n ·Sc.in - Ovu1 ·Sc,lake • (7.2) Evaluating Eq. (7 .2) for the steady state value of the concentration of dissolved ions in the lake yields s;:lake ""104.8 mg kg-I. The difference from the recent mean concentration of dissolved ions indicates that the ion budget of the lake is out of balance. According to 127 Votintsev (1993) the chemical composition of Lake Baikal is unbalanced and the ionic salinity is expected to increase by 3.88 mg kg-I during the next 100 years. Table 7.1: Discharge and ionic salinity of the major inflows and of the only outflow, the Angara. Selenga 31.0 126.8 Upper Angara 8.3 81.3 Barguzin 4.4 134.5 Snezhnaya b)J.7 47.l Turka b)J.S 53.0 Khara-Murin b)1.o 30.2 Utulik b)o.s 64.l total inflows 61.l dl112.o Angara 65.3 e)9s.2 a) Shimaraev et al. (1994) b) M. N. Shimaraev, personal communication c) Votintsev (1993) d) volume weighted mean of ionic salinity of the seven major inflows e) mean ionic salinity in the lake 7.3. Outlook The above example points out one of the main subjects of further research with respect to deep-water ventilation in Lake Baikal. It will be essential to asses the question of how the expected increase in ionic salinity will affect physical and chemical processes in the lake and how it will influence the unique ecosystem in the lake. One way of doing this would be to monitor the concentrations of dissolved chemicals in the major rivers entering the lake. In this way a more detailed model could be obtained for predicting changes in ionic concentrations before they can be detected in the lake. In this thesis, two important sites were identified where deep-water formation occurs due to the delicate interplay between water masses of different temperatures and ionic concentrations. First attempts to quantify these processes were made. However, further efforts are needed to obtain a more consistent picture of deep-water renewal in Lake Baikal. In particular, the following aspects need further research. 128 • Deep-water formation at Academician Ridge was observed but could not be explained in detail. The lateral mixing of water masses with subsequent cabbeling is not yet understood in detail. For the quantification of the deep-water exchange at the sill, further measurements with a better temporal and spatial resolution are needed. For instance, by mooring thermistor chains at the ridge it would be possible to decide whether deep-water formation occurs only during spring and late autumn or whether it endures during the whole winter season. • At this point it is still unclear to which extent deep-water formation at the Academician Ridge is important for the Central Basin. We have shown in Sect. 3.4.2 that during the observation period in spring 1995, the water masses sank to a depth of about 600 mat the southern slope of the sill. It is quite possible that earlier in the year, when surface temperatures near the ridge are lower, colder water masses might penetrate to the deeper part of the Central Basin. Again, it would be helpful to install a thermistor chain at the southern slope of the Academician Ridge. • A rather consistent picture of the deep-water formation occurring at Kukui Canyon was given in Sect. 3.4.1. However, the presented calculations of entrainment factors and of the total water flux to the deepest part of the Central Basin are based on just a few data points. In order to present more reliable data of the importance of the Selenga for deep-water renewal and to be able to apply an entrainment model, more measurements with a better temporal and spatial resolution are required. • It was shown in Sect. 6.4 that in the Southern Basin bottom-water renewal did practically not occur from 1992 to 1995. Yet, the available data do not allow us to decide whether this is part of a long-term saw-tooth fluctuation or whether the basin is shifting towards a permanent state of reduced deep-water renewal. For a better understanding it would be very helpful (1) to continue with the monitoring of the basin with 3ff.3He measurements; (2) to concentrate the investigations with CTD measurements on the southern part of the Selenga Delta where it is likely that river water sinks to the deep part of the Southern Basin; and (3) to supplement the recordings of the thermistor chain in the Southern Basin (Fig. 7 .1) with CTD casts to get information about changes in ionic salinity. • Based on concentrations of CFCs, 3He and 3H, a six-box model of Lake Baikal has recently been developed by Peeters et al. (1997). The model succeeds to reproduce deep- water renewal rates in the three basins. In order to improve the model and to develop a continuous advection-diffusion model, more data are needed, e.g. on the vertical diffusion coefficient K,. References Aeschbach-Hertig, W., 1994. Helium und Tritium als Tracer fiir physikalische Prozesse in Seen. PhD diss. 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Nordic Hydrology 24: 13-30, Appendix Al CTD Profiles: Sampling Positions and Sampling Dates Sampling positions and sampling dates of all CTD profiles taken between 1993 and 1995. 1993 No. Basin Date Loneitude Latitude No. Basin Date Lon2itude Latitude 1 c 16.5.93 106°51'01" 52°41'49" 38 c 28.5.93 107"08'54" 52°48'30" 2 c 18.5.93 107°18'29" 52°38'13" 39 c 28.5.93 107°12'57" 52°44'06" 3 c 18.5.93 107'15'23" 52°40'50" 40 c 28.5.93 107°15'!6" 52°40'43" 4 s 19.5.93 105°16'34" 51°45'05" 41 c 28.5.93 107°16'24" 52°39'1 l" 5 s 21.5.93 105°48'02" 52°13'07" 42 c 28.5.93 107°18'41" 52°38'20" 6 s 21.5.93 106°07'32" 52°29'05" 43 c 28.5.93 107'19'03" 52°36'53" 7 c 21.5.93 106°28'19" 52'25'23" 44 c 29.5.93 106°39'48" 52'34'02" 8 c 21.5.93 106°35'07" 52°30'34" 45 c 29.5.93 106°34'58" 52°28'53" 9 c 21.5.93 106°39'52" 52°33'56" 46 c 29.5.93 106°32'58" 52°27'21" 10 c 23.5.93 106°49'32" 52°40'08" 47 c 29.5.93 !06°28'21" 52°25'19" 11 c 23.5.93 107°18'48" 52'38'25" 48 c 29.5.93 106°27'48" 52'24'55" 12 c 23.5.93 107°16'36" 52°39'12" 49 c 29.5.93 106°41'13" 52'33'15" 13 c 23.5.93 107°15'12" 52°40'37" 50 c 29.5.93 106°46'41" 52'36'50" 14 c 23.5.93 107°13'26" 52°41'56" 51 c 30.5.93 107°13'29" 52°41 '03" 15 c 23.5.93 101•os·o3" 52°47'10" 52 c 30.5.93 107°18'23" 52°38'1 l" 16 c 23.5.93 107'05'01" 52°51'14" 53 c 30.5.93 107°18'30" 52°37'53" 17 c 23.5.93 107°00'22" 52°55'50" 54 c 30.5.93 101°18'15" 52°37'38" 18 c 23.5.93 106°58'13" 52°58'00" 55 c 30.5.93 107'19'12" 52°37'50" 19 c 23.5,93 106°57'18" 52°58'57" 56 c 30.5.93 101°18'56" 52°37'28" 20 c 24.5.93 107°29'01" 52°57'38" 57 c 30.5.93 107'16'20" 52°39'24" 21 c 24.5.93 101°48'20" 53°10'01" 58 c 30.5.93 107°17'00" 52°38'26" 22 c 24.5.93 107°45'45" 53°11'18" 59 c 30.5.93 107'15'23" 52°40'41" 23 c 24.5.93 107°46'37" 53°10'16" 60 c 30.5.93 107°16'35" 52°39'00" 24 c 25.5.93 108°12'37" 53°21'24" 61 c 31.5.93 101°48'03" 53°10'19" 25 c 25.5.93 108°01·40" 53°37'53" 62 c 31.5.93 107°28'58" 52°57'48" 26 N 25.5.93 108°25'09" 53°56'50" 63 c l.6.93 106°48'44" 53°39'36" 27 N 25.5.93 108°46'20" 54°18'49" 64 s 1.6.93 105°56'54" 52°05'58" 28 N 25.5.93 108°55'14" 54°28'46" 65 s 1.6.93 105°58'40" 52°03'00" 29 N 26.5.93 108°13'28" 53°48'28" 66 s 1.6.93 105°54'09" 52°00'43" 30 c 26.5.93 108°01'43" 53°43'36" 67 s l.6.93 105°16'43" 51°44'36" 31 c 26.5.93 107°45'45" 53°37'00" 68 s 3.6.93 104°54'42" 51°49'36" 32 c 26.5.93 107°42'39" 53°34'41" 69 c 5.6.93 107°18'32" 52°38'18" 33 c 26.5.93 107°37'39" 53°26'02" 70 c 5.6.93 107°18'35" 52°37'54" 34 c 27.5.93 106°57'18" 52°58'22" 71 c 5.6.93 107°18'14" 52°36'36" 35 c 27.5.93 106°58'04" 52°57'42" 72 c 5.6.93 107°18'20" 52°37'21" 36 c 28.5.93 107°00'05" 52°56'02" 73 c 5.6.93 107°17'05" 52°37'13" 37 c 28.5.93 107°04'57" 52°52'03" 74 c 5.6.93 107'17'29" 52°38'14" 136 No. Basin Date Lon111tude Latitude No. Basin Date Lon1dtude Latitude 75 c 5.6.93 107°16'16" 52°39'04" 138 c 16.6.93 107°18'49" 52°37'22" 76 c 5.6.93 !07°15'42" 52°39'52" 139 c 16.6.93 !07°18'46" 52°36'40" 77 c 5.6.93 107°15'08" 52°40'53" 140 N 17.6.93 !08°25'18" 53°56'27" 78 c 6.6.93 107°13'45" 52°41'36" 141 N 17.6.93 109'03'52" 54°27'16" 79 c 7.6.93 107°19'04" 52'36'47" 142 N 18.6.93 109°28'44" 55°19'20" 80 c 7.6.93 107°18'34" 52°37'18" 143 N 18.6.93 109°51'58" 55°41'46" 81 c 7.6.93 107°18'20" 52°38'04" 144 N 18.6.93 109°51'09" 55°40'52" 82 c 7.6.93 107'18'19" 52°38'40" 145 N 18.6.93 109°50'33" 55°39'57" 84 c 7.6.93 107'17'1 I" 52°39'27" 146 N 18.6.93 109°49'54" 55°38'57" 85 c 7.6.93 107°15'44" 52°39'43" 147 N 18.6.93 109°48'54" 55°38'09" 86 c 7.6.93 107°15'20" 52°40'40" 148 N 18.6.93 109'50'33" 55°39'28" 87 c 7.6.93 107°15'09" 52°40'47" 149 N 18.6.93 109°48'49" 55°31 '07" 88 c 7.6.93 107°13'14" 52°41'51" 150 N 18.6.93 109°47'53" 55'31'12" 89 c 7.6.93 107°14'09" 52°41'02" 151 N 18.6.93 109'47'03" 55°31'02" 90 c 8.6.93 107°18'33" 52°38'22" 152 N 18.6.93 109°46'09" 55°30'59" 91 c 8.6.93 107°15'13" 52°40'50" 153 N 18.6.93 109°44'52" 55°31'05" 92 c 8.6.93 107°13'52" 52°42'15" 154 N 18.6.93 109'43'11" 55°30'54" 93 c 8.6.93 107°10'53" 52°44'52" 155 N 19.6.93 109°34'43" 55°34'10" 94 c 8.6.93 107°07'41" 52°48'00" 156 N 19.6.93 109°31'10" 55°25'23" 95 c 8.6.93 107°05'28" 52°50'34" 157 N 19.6.93 109'25'48" 55°02'30" 96 c 8.6.93 107°00'01" 52°56'01" 158 N 19.6.93 108°46'34" 54°18'44" 97 c 8.6.93 107°57'55" 52°57'49" 159 N 19.6.93 108°46'34" 54'18'44" 98 c 8.6.93 106'57'04" 52°58'46" 160 N 20.6.93 108°25'29" 53°56'23" 99 c 9.6.93 106°53'00" 53°04'18" 161 N 20.6.93 109°07'39" 53°49'58" 100 c 9.6.93 107°08'19" 53°11'45" 162 N 20.6.93 109°08'00" 53°51 '02" IOI c 9.6.93 106°23'54" 53°18'37" 163 N 20.6.93 109°08'44" 53°51'51" 102 c 9.6.93 107°39'58" 53°26'30" 164 N 20.6.93 !09°08'53" 53°52'32" 103 c 10.6.93 107°47'48" 53°10'56" 165 N 20.6.93 109°09'25" 53°55'45" 104 c 10.6.93 108°09'27" 52°56'00" 166 N 20.6.93 108°49'55" 53°52'22" 105 c 10.6.93 108°05'00" 52°55"10" 168 c 20.6.93 108°44'08" 53°48'48" 106 c 10.6.93 107°52'27" 52'56'30" 169 c 20.6.93 108°07'56" 53'37'57" 107 c 10.6.93 107°54'19" 52°57'25" 170 c 20.6.93 108°03'46" 53°39'34" 108 c 10.6.93 107'05'38" 52°35'37" 171 c 20.6.93 108°06'34" 53°31'51" 109 c 10.6.93 107°06'13" 52°36'50" 172 c 20.6.93 108°13'00" 53°21'15" l 10 c 11.6.93 107°19'1 I" 52'36'54" 173 c 21.6.93 107°43'03" 54°11 '22" Ill c 11.6.93 101°18·22" 52'37'15" 174 c 22.6.93 107°28'58" 52°57'44" 112 c 11.6. 93 107°19'07" 52°37'05" 175 c 22.6.93 106°57'3 I" 52°58'46" 113 c 11.6.93 107°19'22" 52°37'12" 176 c 22.6.93 106°58'14" 52°57'45" 114 c 11.6.93 107°19'07" 52°37'41" 177 c 24.6.93 107°18'53" 52°36'40" 115 c 11.6.93 107°17'21" 52°38'13" 178 c 24.6.93 107°18'29" 52°37'07" 116 c 11.6.93 107°17'14" 52°38'48" 179 c 24.6.93 107°18'05" 52°37'33" 117 c 11.6.93 107°16'54" 52°39'33" 180 c 24.6.93 107°17'33" 52°38'07" 118 c 11.6.93 107°15'46" 52°39'43" 181 c 24.6.93 107°16'26" 52°39'00" 119 c 1 l.6.93 101°15'36" 52°40'49" 182 c 24.6.93 101°15'18" 52°40'30" 120 11.6.93 183 c 24.6.93 107°14'55" 52°40'09" 121 c 12.6.93 107°18'39" 52'38'17" 184 c 24.6.93 107°13'43" 52'42'01" 122 c 12.6.93 107°15'28" 52°39'42" 185 c 24.6.93 107°10'31" 52°44'44" 123 c 12.6.93 107°15'18" 52°40'44" 186 c 24.6.93 107°07'41" 52°48'04" 124 c 12.6.93 107°13'47" 52°42'07" 187 c 24.6.93 107'04'03" 52°52'05" 126 c 12.6.93 107°07'46" 52°47'36" 188 c 24.6.93 107°00'01" 52°55'56" 127 c 12.6.93 107°04'35" 52°s1•14" 189 c 24.6.93 106°58'02" 52°57'55" 128 c 12.6:93 106°59'48" 52°55'57" 190 c 24.6.93 106°57'05" 52°58'51" 129 c 12.6.93 106°57'01" s2°5B'54" 191 c 25.6.93 106°49'04" 52'39'39" 130 c 12.6.93 106°58'01" 52°57'46" 192 c 25.6.93 106'39'52" 52'33'55" 131 c 16.6.93 107°13'14" 52°42'02" 193 c 25.6.93 106°28'08" 52°25·25" 132 c 16.6.93 107°14'42" 52'41'18" 194 c 25.6.93 106°07'31" 52°29'06" 133 c 16.6.93 107°15'19" 52°40'53" 195 s 25.6.93 105°48'20" 52°13'32" 134 c 16.6.93 107°16'08" 52°39'59" 196 s 26.6.93 105°16'58" 51°45'16" 135 c 16.6.93 107°16'49" 52°39'11" 197 s 26.6.93 104'36'56" 51°40'58" 136 c 16.6.93 107°18'04" 52'38'18" 198 s 26.6.93 104°13'1 l" 51°41'27" 137 c 16.6.93 107°17'41" 52°37'31" 137 1994 No. Basin Date Loneitude Latitude No. Basin Date Loneitude Latitude I s 21.10.94 51°40'51" 103°54'47" 64 s 4.11.94 51°38'34" 104°13'42" 2 s 21.10.94 51°45'1 l" 104°12'41" 65 s 4.11.94 51°33'39" 104°14'01" 3 s 21.10.94 51°44'07" 104°13'13" 66 s 4.11.94 51°32'39" 104°14'15" 4 s 21.10.94 51°43'03" 104°11'52" 67 s 4.11.94 51°31 '19" 104°14'11" 5 s 22.10.94 51°38'33" 104°13'27" 68 s 4.11.94 51°41'04" 104°37'13" 6 s 22.10.94 51°33'32" 104°14'14" 69 s 5.11.94 51°41'30" 105°02'03" 7 s 22.10.94 5 I 0 32'45" 104°14'18" 70 c 5.11.94 51°57'04" 107°30'22" 8 s 22.10.94 51°31 '45" 104°14'25" 71 c 5.11.94 53°11'08" 107°46'02" 9 s 22.10.94 51°40'51" 104°36'54" 72 c 6. I 1.94 53°22'17" 108°12'40" IO s 22.10.94 51°41'44" 105°00'34" 73 c 6.11.94 53°37'43" 108°07'34" 11 s 23.10.94 51°44'45" 105°16'58" 74 N 6. I 1.94 53°56'40" 108°25'22" 12 s 23.10.94 52°13'45" 105°48'17" 75 N 7.11.94 54°19'08" 108°46'09' 13 s 23.10.94 52°28'22" 106°07'17" 76 N 7.11.94 54°29'05" 108°52'04" 14 c 23.10.94 52°24'39" 106°27'40" 77 N 7.11.94 54°36'50" 108°56'37" 15 c 23.10.94 52°29'59" 106°36'00" 78 N 7.11.94 55°03'07" 109°15'19" 16 c 23.10.94 52°33'53" 106°39'34" 79 N 7.11.94 55°34'26" 109°35'08" I 7 c 23. I 0.94 52°39'32" 106°49'16" 80 N 8.11.94 55°43'30" 109°36'32" 18 c 24.10.94 52°37'30" 107°17'51" 81 N 8.11.94 55°24'15" 109°29'34" 19 c 24.10.94 52°39'02" 107°16'24" 82 N 8.11.94 55°29'53" 109°31'05" 20 c 24.10.94 52°40'34" 107°15'18" 83 N 8.11.94 55°29'37" 109°31 '38" 21 c 24.10.94 52°4 I '55" 107°13'40" 84 N 8.11.94 55°29'14" 109°32'05" 22 c 24.10.94 52°48'03" 107°07'50" 85 N 8.11.94 55°29'15" 109°32'33" 23 c 24.10.94 52°52'05" 107°04'08" 86 N 8.11.94 55°29'03" 109°32'57" 24 c 24. I 0.94 52°56'00" 107°00'24" 87 N 8.11.94 55°28'08" 109°32'19" 25 c 24. I 0.94 52°59'02" 106°57'04" 88 N 8.11.94 55°27'15" 109°31 '48" 26 c 24.10.94 52°57'47" 107°29'29" 89 N 8.11.94 55°26'57" 109°32'13" 27 c 24.10.94 53°11'07" 107°45'52" 90 N 8.11.94 55°25'15" 109°30'10" 28 c 25.10.94 53°37'55" 108°07'59" 91 N 8.11.94 55°26'11" 109°30'53" 29 N 25.10.94 53°56'48" 108°25'20" 92 N 8.11.94 55°26'51" 109°31 '27" 30 N 25.10.94 54°19'11" 108°46'11' 93 N 9.11.94 55°21 '09" 109°13'51" 31 N 25.10.94 55°03'05" 109' I 4'57' 94 N 9.11.94 55°21'17" 109°16'11" 32 N 26.10.94 55°17'43" 109°42'46' 95 N 9.11.94 55°20'59" 109°18'05" 33 N 26.10.94 55°17'50" 109°40'40' 96 N 9.11.94 55°20'43" 109°21'05" 34 N 26.10.94 55°18'09" 109°38'26' 97 N 9.11.94 55°19'57" 109°24'25" 35 N 26.10.94 55°18'49" 109°34'13' 98 N 9.11.94 55°19'33" 109°29'36" 36 N 26.10.94 55°19'33" 109°29'16' 99 N 9.11.94 55°18'47" 109°34'26" 37 N 26.10.94 55°20'01" 109°24'16' 100 N 9.11.94 55°18'07" 109°38'37" 38 N 26.10.94 55°20'42" 109°19'33" IOI N 9.11.94 55°17'50" 109°41'06" 39 N 26.10.94 55°21'12" 109°17'32" 102 N 9.11.94 55°17'41" 109°42'51" 40 N 26.10.94 55°21'33" 109°15'56" 103 N 9.11.94 55°17'28" 109°44'36" 41 N 27.10.94 55°34'25" 109°35'05" 104 c 9.11.94 53°10'49" 107°47'15" 42 N 28.10.94 55°45'53" 109°36'55" 105 c 11.11.94 52°59'03" 106°57'20" 43 N 28.10.94 55°45'50" 109°36'54" 106 c 11.11.94 52°57'39" 106°58'14" 44 N 28.10.94 55°45'42" 109°36'43" 107 c 11.11.94 52°56'08" 107°00'12" 45 N 28.10.94 55°45'42" 109°36'39" 108 c 11.11.94 52°52'52" 107°04'04" 46 N 28.10.94 55°45'40" 109°36'35" 109 c 11.11.94 52°48'03" 107°08'22" 47 N 28.10.94 55°45'36" 109°36'28" 110 c 11.11.94 52°44'44" 107°11'25" 48 N 28.10.94 55°45'35" 109°36'19" Ill c 11.11.94 52°41'43" 107°15'08" 49 N 28.10.94 55°45'32" 109°36'14" 112 c 11.11.94 52°40'37" 107°15'24" 50 N 28.10.94 55°45'30" 109°36'07" 113 c 11.11.94 52°39'00" 107°16'40" 51 N 28.10.94 55°45'34" 109°36'05" 114 c 11.11.94 52°37'34" 107°17'43" 52 N 28.10.94 55°45'38" 109°36'05" 115 c 12.11.94 52°39'50" 106°49'49" 53 N 28.10.94 55°45'13" 109°36'07" 116 c 12.11.94 52°33'56" 106°40'12" 54 N 28.10.94 55°45'06" 109°36'00" 117 c 12.11.94 52°29'49" 106°35'26" 55 N 28.10.94 55°43'29" 109°36'42" 118 c 12.11.94 52°26'26" 106°31'50" 56 N 29.10.94 54°37'03" 108°57'20" 119 c 12.11.94 52°24'47" 106°27'59" 57 N 29.10.94 54°29'41" 108°52'22" 120 c 12.11.94 52°28'21" 106°06'50" 58 c 31.10.94 53°21 '59" 108°12'36" 121 s 12.11.94 52°23'40" 105°59'25" 59 s 4.11.94 51°41'05" 103°54'59" 122 s 12.11.94 51°16'41" 105°16'41" 60 s 4.11.94 51°46'38" 104°14'09" 123 s 14.11.94 51°46'29" 104°13'21" 61 s 4.11.94 51°44'35" 104°13'55" 124 s 14.11.94 51°45'25" 104°12'43" 62 s 4.11.94 51°45'34" 104°14'04" 125 s 14.11.94 51°43'56" 104°11 '26" 63 s 4.11.94 51°43'35" 104°13'14" 126 s 14.11.94 51°43'06" 104°12'06" 138 1995 No. Basin Date Lon1dtude Latitude No. Basin Date Lonuitude Latitude 1 s 11.5.95 51'41'59" 104'58'51" 65 c 17.5.95 52°56'17" 107°00'35" 2 s 12.5.95 51°35'21" 105°07'16" 66 c 17.5.95 52°57'59" 106°58'01" 3 s 12.5.95 51°30'19" 104°14'38" 67 c 17.5.95 52'59'04" 106'57'09" 4 s 12.5.95 51°31'03" 104°14'47" 68 c 17.5.95 52'39'38" 106°49'13" 5 s 12.5.95 51°3149" 104°14'02" 69 c 18.5.95 52°33'43" 106°39'32" 6 s 12.5.95 51°33'21" 104°13'25" 70 c 18.5.95 52°30'59" 106°37'20" 7 s 12.5.95 51°34'17" 104°14'04" 71 c 18.5.95 52°26'57" 106°32'03" 8 s 12.5.95 51°34'19" 104°16'26" 72 c 18.5.95 52°24'39" 106°27'28" 9 s 12.5.95 51'32'23" 104°16'02" 73 c 18.5.95 52°44'40" 105°17'00" 10 s 12.5.95 51°31'38" 104'15'52" 74 c 20.5.95 51°54'13" 105°05'27" 11 s 12.5.95 51'32'00" 104°15'33" 75 c 20.5.95 51°54'05" 105°05'27" 12 s 12.5.95 51°33'??" 104'14'38" 76 c 20.5.95 51°53'55" 105°05'27" 13 s 13.5.95 51°31'57" 104°09'39" 77 c 20.5.95 51°53'40" 105'05'27" 14 s 13.5.95 51°33'24" 104°09'37" 78 c 20.5.95 51'53'28" 105'05'06" 15 s 13.5.95 51°34'30" 104'10'29" 79 c 20.5.95 51°52'59" 105°05'16" 16 s 13.5.95 51°35'42" 104°10'52" 80 c 20.5.95 51°52'43" 105°05'26" 17 s 13.5.95 51°37'41" 104°11'35" 81 c 20.5.95 51 '44'57" 105°17'07" 18 s 13.5.95 51°39'32" 104'12'26" 82 c 2!.5.95 52°15'12" !05°43'55" 19 s 13.5.95 51°43'26" 104°12'34" 83 c 21.5.95 52°14'40" 105°43'42" 20 s 13.5.95 51°44'30" 104°12'50" 84 c 21.5.95 52°13'54" 105°43'47" 21 s 13.5.95 51°45'40" 104°13'00" 85 c 21.5.95 52°13'21" 105°43'50" 22 s 13.5.95 51°46'23" 104°12'54" 86 c 21.5.95 52'28'52" 106°40'41" 23 s 13.5.95 51°46'57" 104°13'05" 87 c 21.5.95 52°30'45" 106°38'37" 24 s 13.5.95 51°46'06" 104°12'46" 88 c 21.5.95 52°32'28" 106°38'07" 25 s 13.5.95 51°45'58" 104'08'57" 89 c 21.5.95 52°36'16" 106°36'32" 26 s 13.5.95 51°45'47" 104°08'32" 90 c 21.5.95 52°34'48" 107°06'01" 27 s 13.5.95 51°45'16" 104°08'33" 91 c 21.5.95 52°34'40" 107°05'50" 28 s 13.5.95 51°40'57" 103'53'53" 92 c 21.5.95 52'34'40" 107°05'48" 29 s 13.5.95 51°43'53" 103°56'44" 93 c 2 l.5.95 52°34'38" 107°05'44" 30 s 13.5.95 51°44'21" 103°56'49" 94 c 21.5.95 52°34'34" 107°05'40" 31 s 14.5.95 51°44'27" 103°56'50" 95 c 21.5.95 52°34'35" 107°05'39" 32 s 14.5.95 51°46'30" 104°25'34" 96 c 21.5.95 52'34'21" 107°05'44" 33 s 14.5.95 51°46'54" 104°25'03" 97 c 21.5.95 52°35'55" 107°05' 17" 34 s 14.5.95 51°47'12" 104°26'11" 98 c 21.5.95 52°36'5 l" 107°04'39" 35 s 14.5.95 51°47'29" 104°26'23" 99 c 21.5.95 52°38'00" 107°03'45" 36 s 14.5.95 51°38'05" 104°35'29" 100 c 22.5.95 52°58'09" 107°57'1 l" 37 s 14.5.95 51°41'58" 105'00'32" IOI c 22.5.95 52°58'34" 107°57'00" 38 s 14.5.95 51°44'58" 105°16'53" 102 c 22.5.95 53°11'00" 107°45'43" 39 s 15.5.95 52°15'05" 105°43'19" 103 c 22.5.95 53°12'51 .. 107°44'22" 40 s 15.5.95 52°15'05" 105°44'21" 104 c 22.5.95 53°13'21" 107°44'15" 41 s 15.5.95 52°14'46" 105°44'57" 105 c 22.5.95 53°13'31" 107'44'00" 42 s 15.5.95 52°13'23" 105°47'49" 106 c 22.5.95 53'21 '53" 108'13'05" 43 c 15.5.95 52°28'37" 106°06'56" 107 c 22.5.95 53°23'21" 108°38'56" 44 c 15.5.95 52°24'39" 106'27'28" 108 c 22.5.95 53°23'38" 108°40'56" 45 c 15.5.95 52°26'57" 106°32'02" 109 c 22.5.95 53°24'02" 108°44'33" 46 c 15.5.95 52'30'11" 106'34A57" 110 c 22.5.95 53°26'27" 108°39'18" 47 c 15.5.95 52°34'19" 106°40'50" 111 c 22.5.95 53°28'04" 108°35'37" 48 c 15.5.95 52°39'32" 106°49'12" 112 c 22.5.95 53°28'45" 108'34'18" 49 c 15.5.95 52°25'54" 106°24'22" 113 c 22.5.95 53°29'40" 108°33'04" 50 c 16.5.95 52°49'25" 106°43'36" 114 N 22.5.95 53°38'30" 108°08' 17" 51 c 16.5.95 52°49'23" 106°43'32" 115 N 23.5.95 53°56'12" 108°25'18" 52 c 16.5.95 52°49'09" 106°45'44" 116 N 23.5.95 53°54'33" 108°29'29" 53 c 16.5.95 53°11 '06" 107°45'40" 117 N 23.5,95 53°58'13" 108°22'04" 54 c 16.5.95 53°11'16" 107°46'15" 118 N 23.5.95 54°00'49" 108°16'54" 55 c 16.5.95 52°57'58" 107°29'59" I 19 N 23.5.95 53°53'37" 108°32'45" 56 c 17.5.95 52'36'31" 107°18'25" 120 M 23.5.95 53°45'48" 108°36'49" 57 c 17.5.95 52°36'58" 107°17'52" 121 M 23.5.95 53°49'53" 108°43'30" 58 c 17.5.95 52°37'35" 107°17'42" 122 M 23.5.95 53°33'29" 108'18'57" 59 c 17.5.95 52°39'04" 107°16'21" 123 M 23.5.95 53°35'50" 108'12'38" 60 c 17.5.95 52°40'32" 107°15'00" 124 N 23.5.95 53°39'06" 108'04'09" 61 c 17.5.95 52°42'03" 107°13'55" 125 N 23.5.95 53°40'01" 108'01'50" 62 c 17.5.95 52°42.13" 107'14'28" 126 N 24.5.95 54°07'35" 108°35'35" 63 c 17.5.95 52"48'23" 107'08'54" 127 N 24.5.95 54°19'17" 108°46'22" 64 c 17.5.95 52°51'42" 107'05'57" 128 N 24.5.95 54°29'21" 108°53'45" 139 No. Basin Date Longitude Latitude No. Basin Date Loneilude Latitude 129 N 24.5.95 54°30'02" 108°57'3 l" 195 M 30.5.95 52°47'54" 107°07'48" 130 N 24.5.95 54°30'46" 108°46'54" 196 M 30.5.95 52°44'16" 107°10'54" 131 N 24.5.95 54°31'31" 108°42'37" 197 M 30.5.95 52°42'05" 107°13'56" 132 M 24.5.95 54°31 '56" 108°40'27" 198 M 30.5.95 52°40'31" 107°14'58" 133 M 24.5.95 54°32'07" 108°40'15" 199 M 30.5.95 52°39'02" 107°16'15" 134 N 24.5.95 54°32'13" 108°39'59" 200 M 30.5.95 52°37'27'' 107°17'43" 135 N 24.5.95 54°35'01" 108°54'27" 201 M 30.5.95 52°36'52" 107°17'54" 136 N 25.5.95 55°02'09" 109°06'58" 202 M 30.5.95 52°36'30" 101°18'04" 137 N 25.5.95 55°01 '55" 109°07'1 l" 203 M 31.5.95 52°34'56" 107°05'59" 138 N 25.5.95 55°01'50" 109'07'36" 204 M 31.5.95 52°35'51" 107°04'54" 139 N 25.5.95 55°01 '36" 109°09'33" 205 M 31.5.95 52°36'45" 107°04'42" 140 N 25.5.95 55°00'39" 109'12'33" 206 M 31.5. 95 52°37'59" 107°03'49" 141 N 25.5.95 54°58'31" 109°20'45" 207 M 31.5.95 52°39'3 I" 106°46'08" 142 N 25.5.95 54°56'22" 109°28'50" 208 M 31.5.95 52°34'12" 106°40'52" 143 N 25.5.95 54°54'04" 109°37'00" 209 M 31.5.95 52°30'07'' 106°34'59" 144 N 25.5.95 54°56'12" 109°33'54" 210 M 31.5.95 52°26'54" 106°31'54" 145 N 25.5.95 55°06'09" 109°39'05" 211 M 31.5.95 52°24'47" !06°27'21" 146 N 25.5.95 55°06'22" 109°39'17" 212 M 31.5.95 52°27'19" 106°10'!1" 147 N 25.5.95 55°!7'15" !09°42'50" 213 M 3 l.5.95 52°28'43" 106°06'50" 148 N 25.5.95 55°17'45" 109°41'30" 214 s 31.5.95 52°13'15" !05°47'56" 149 N 25.5.95 55°19'12" 109°39'03" 215 s 1.6.95 52°01'56" 106°00'01" 150 N 25.5.95 55°18'42" 109°33'51" 216 s 1.6.95 52°00'00" 105°54'01" 151 N 25.5.95 55°19'19" 109°29'04" 217 s 1.6.95 51°57'38" 105°45'1 I" 152 N 25.5.95 55°20'14" 109°24'04" 218 s 1.6.95 51°54'37" 105°36'49" 153 N 26.5.95 55°21'16" 109°41'21" 219 s 1.6.95 51°44'57" 105°16'50" 154 N 26.5.95 55°21'13" 109'16'01" 220 s 1.6.95 51°42'15" 105'00'33" 155 N 26.5.95 55°21'13" 109°16'01" 221 s 1.6.95 51°38'00" 104°37'10" 156 N 26.5.95 55°43'1 I" 109'36'44" 222 s 2.6.95 51°47'26" 104°26'23" 157 N 26.5.95 55"34'04" 109°34'58" 223 s 2.6.95 51°47'39" 104°26'19" 158 N 27.5.95 55°17'25" 109°12'54" 224 s 2.6.95 51°47'10" 104°26'04" 159 N 27.5.95 55°16'59" 109°13'28" 225 s 2.6.95 51'46'54" 104'25'00" 160 N 27.5.95 55°16'23" 109°13'59" 226 s 2.6.95 51°46'40" 104°25'36" 161 N 27.5.95 55°00'35" 109°11 '42" 227 s 2.6.95 51 '40'54" 103°53'46" 162 N 27.5.95 54'59'03" 109°10'10" 228 s 2.6.95 51°42'07" 103°48'17'' 163 N 27.5.95 54°58'53" 109°09'47" 229 s 2.6.95 51°42'40" 103°45'38" 164 N 27.5.95 54°34'42" 108°54'24" 230 s 2.6.95 51°42'45" 103°45'06" 165 N 28.5.95 54°32'02" 1os0 40·21" 231 s 2.6.95 51°43'1 l" 103°44'23" 166 N 28.5.95 54°32'14" 108°40'05" 232 s 2.6.95 51°43'23" 103°43'51" 167 N 28.5.95 54'31 '24" 108°42'52" 233 s 2.6.95 52°36'40" 103°54'49" 168 N 28.5.95 54°29'25" 108°53'58" 234 s 2.6.95 52°36'58" 103°54'59" 169 N 28.5.95 54°27'04" 109°04'12" 235 s 2.6.95 52'36'49" 103°55'05" 170 N 28.5.95 54°25'52" 109°08'52" 236 s 2.6.95 52°36'42" 103°55'1 I" 171 N 28.5.95 54°23'49" 109°16'44" 237 s 2.6.95 52°36'44" 103°55'1 l" 172 N 28.5.95 54°23'04" 109°20'59" 238 s 2.6.95 52°36'42" 103'55'11" 173 N 28.5.95 54'22'51" 109°22'2[" 239 s 2.6.95 52'36'38" 103°55'10" 174 N 28.5.95 54°18'23" 108°43'56" 240 s 3.6.95 52°36'38" 103°55'08" 175 N 28.5.95 53'56'21" 108'24'58" 241 s 3.6.95 52'37'09" 103°55'14" 176 N 29.5.95 53°40'00" 108°01'54" 242 s 3.6.95 52°37'35" 103°55'23" 177 N 29.5.95 53°39'00" 1os0 04'13" 243 s 3.6.95 52°32'02" 104°09'42" 178 M 29.5.95 53°38'26" 108°08'15" 244 s 3.6.95 52°32·21" 104°09'55" 179 M 29.5.95 53°35'42" 108°12'34" 245 s 3.6.95 52°33'29" 104°10'15" 180 M 29.5.95 53'33'27" 108°19'07" 246 s 3.6.95 52°34'31" 104°10'33" 181 M 29.5.95 53°31'27" 108°23'38" 247 s 3.6.95 52'35'34" 104°10'50" 182 M 29.5.95 53°31'10" 108°27'01" 248 s 3.6.95 52°37'39" 104°1 l '03" 183 M 29.5.95 53°29'37" 108°29'33" 249 s 3.6.95 52'39'33" 104°11'30" 184 M 29.5.95 53°21 '44" 108°13'!3" 250 s 3.6.95 52°41 '13" 104°1 l '53" 185 M 29.5.95 53°11'05" 107°45'34" 251 s 3.6.95 52°43'25" 104°12'34" 186 M 29.5.95 52'57'41" 107°29'39" 252 s 3.6.95 52°44'27" 104°12'53" 187 M 29.5.95 53°02'18" I07'25'27" 253 s 3.6.95 52°45'42" 104°12'45" 188 M 29.5.95 53°03'48" 107°24'32" 254 s 3.6.95 52°46'04" 104°12'47" 189 M 29.5.95 53°04'06" 107°24'19" 255 s 3.6.95 52°46'26" 104°12'52" 190 M 30.5.95 52°58'04" 106°58'07" 256 s 3.6.95 52°46'53" 104'13'01" 191 M 30.5.95 52°58'54" 106°57'1 l" 257 s 4.6.95 52°48'00" 104°55'22" 192 M 30.5.95 52°59'10 106°56'47" 258 s 4.6.95 52°41'40" 104'58'49" 193 M 30.5.95 52'56'!3" 107°00'04" 259 s 4.6.95 52'36'53" 104'01'18" 194 M 30.5.95 52°51'18" 107°04'21" 140 A2 Results of Noble Gas and Tritium Analysis The measured 4He, 20Ne and 3H concentrations and the 3He/4He ratios are summarised in the table below. Errors ar given as lcr values. Positions SI, Cl and Nl are the deepest positions of each basin; positions S2, C2, C3, C4 and N2 are shown in Fig. 6.1; FB, Kot, Ksy and FB are hydrothermal springs (see Fig. 6.1); Kl - K4 are positions in Kukui Canyon at water depths of about 200 m, 400 m, 600 m, 800 m and 1200 m, respectively; positions C5 (53°22'17" N; 108°12'40" E) and N3 (53°56'27" N; 108°25'20" E) are two additional stations in the Central and Northern Basins. Samples that were not taken at one of these positions are not included in the table. pos date depth T "H 4 He ilffe/4 He 211Ne [ml [•CJ [TU] no-8 cm3sTP .-11 no-61 no-7 cm3STPo·l SI 23.3.92 0 0.5 21.6±0.6 4.671±0.029 J.390±0.008 l. 977±0.016 SI 23.3.92 50 l 19.4±0.5 4.698±0.029 1.455±0.008 1.889±0.015 SI 26.3.92 150 3 18.6±0.5 4.932±0.030 1.569±0.008 1.961±0.015 SI 26.3.92 300 3.5 18.4±0.6 4.658±0.029 1.656±0.011 1.866±0.014 SI 26.3.92 400 3.5 18.3±0.6 4.615±0.028 1.719±0.013 1.852±0.015 SI 26.3.92 460 3.5 18.3±0.5 4.640±0.029 1.782±0.014 1.854±0.015 SI 26.3.92 700 3.4 16.7±0.5 4.690±0.029 1.986±0.017 1.872±0.015 SI 26.3.92 900 3.4 16.1±0.5 4.691±0.030 2.139±0.022 l.862±0.014 SI 26.3.92 llOO 3.4 16.0±0.6 4.706±0.030 2.329±0.025 1.853±0.014 SI 26.3.92 1350 3.4 l 5.9±0.4 4. 763±0.032 2.220±0.030 1.866±0.015 Cl 14.6.92 0 3.5 19.3±0.6 4.705±0.029 1.444±0.008 1.879±0.014 Cl 14.6.92 100 3.5 20.3±0.5 4.654±0.028 1.495±0.008 1.882±0.014 Cl 14.6.92 300 3.5 20.2±0.6 4.664±0.029 1.623±0.01 l 1.852±0,016 Cl 14.6.92 500 3.4 19.3±11.0 4.685±0.029 J.740±0.025 1.863±0.015 Cl 14.6.92 700 3.3 18.3±0.l 5.760±0.036 1.820±0.017 1.985±0.016 Cl 14.6.92 930 3.3 17.2±1.5 4.814±0,031 2.259±0.023 1.871±0.015 Cl 14.6.92 1200 3.2 15.2±0.5 4. 772±0.031 2.488±0.028 1.865±0.016 Cl 14.6.92 1400 3.2 14.1±0.5 4.841±0.032 2.559±0.032 1.866±0.014 CJ 14.6.92 1650 3.2 15.2±0.5 5.225±0.035 2.190±0.036 1.908±0.014 C2 15.6.92 300 3.5 0.0±0.0 0.000±0.003 0.000±0.006 0.000±0.000 C2 15.6.92 580 3.4 19.6±0.6 4. 725±0.029 1.844±0.015 l.866±0.015 C2 15.6.92 900 3.3 17.1±0.5 4. 793±0.030 2.316±0.021 1.856±0.014 C2 15.6.92 1100 3.25 15.1±0.7 4.813±0.031 2.480±0.026 1.866±0.017 C2 15.6.92 1300 3.2 13.7±0.5 4.697±0.031 2.501±0.029 1.835±0.014 FB 7.7.92 396 3.5 19.3±0.5 6.058±0.043 1.504±0.011 1.829±0.019 FB 7.7.92 422 3.5 18.8±0.5 5.110±0.031 1.836±0.013 1.856±0.014 FB 7.7.92 430 3.5 19.9±0.6 5.139±0.032 1.799±0.012 J.846±0.015 FB 7.7.92 430 3.5 20.2±0.5 5.129±0.032 1.781±0.012 1.876±0.018 FB 7.7.92 445 3.5 19.5±0.4 7.502±0.046 J.290±0.012 1.830±0.014 FB 8.7.92 463 3.5 18.9±0.5 11.868±0.084 0.895±0.010 1.851±0.020 FB 9.7.92 425 3.5 0.0±0.0 5.216±0.032 1.807±0.014 1.871±0.016 Ksy 9.7.92 0 46 9.3±0.4 790.495±4.877 O. l 79±0.001 1.676±0.016 Ksy 9.7.92 0 46 9.3±0.4 799.223±5.712 0.181±0.001 1.699±0.018 PB 11.7.92 453 3.5 19.2±0.6 8.558±0.061 1.179±0.013 1.845±0.019 FB 11.7.92 453 3.5 20.0±0.6 6.859±0.042 1.386±0.012 l.873±0.018 FB 11.7.92 670 3.5 0.0±0.l 5.118±0.032 2.015±0.016 1.848±0.017 NI 12.7.92 100 3.5 20.0±0.5 4. 734±0.029 1.498±0.009 1.869±0.0I5 NI 12.7.92 240 3.5 20.5±0.6 4. 780±0.029 l.512±0.009 1.870±0.018 NI 12.7.92 400 3.5 20.0±0.4 4.794±0.029 1.669±0.012 1.835±0.015 NI 12.7 .92 600 3.45 I 9.3±0.5 5.124±0.032 1.865±0.016 1.861±0.015 NI 12.7.92 800 3.4 19.0±0.5 5.327±0.033 l.953±0.019 1.856±0.014 NI 12.7.92 880 3.4 18.8±0.5 5.366±0.034 1.834±0.020 1.867±0.015 SI 14.7.92 50 5 18.4±0.5 4.614±0.028 l.533±0.009 1.851±0.019 SI I4.7.92 150 3.8 18.8±0.6 4.589±0.028 1.563±0.009 0.000±0.000 SI 14.7.92 300 3.6 18.4±0.5 4.651±0.029 1.619±0.011 J.871±0.014 141 pos date depth T 3H 4He 3He/4He ZUNe [m] [OC] [TIJ] r10-8 cm3sTP ,-11 r10-61 no-7 cm3sTPe-l' SI 14.7.92 500 3.5 17.7±0.5 4. 732±0.029 1.912±0.014 1.867±0.016 SI 14.7.92 700 3.4 15.5±0.5 4.946±0.031 2.137±0.018 1.877±0.014 SI 14.7.92 900 3.4 14.6±0.5 6.163±0.038 2.113±0.021 2.193±0.017 SI 14.7.92 1100 3.35 14.6±0.5 5.021±0.032 2.286±0.025 1.861±0.015 SI 14.7.92 1300 3.35 14.6±0.4 5.058±0.033 2.251±0.029 1.876±0.015 SI 19.5.93 0 2.6 18.8±0.5 4.634±0.017 1.436±0.008 1.883±0.013 SI 19.5.93 600 3.4 17.9±1.3 24.168±0.092 1.500±0.015 SI 19.5.93 1000 3.4 15.3±1.3 4.852±0.020 2.197±0.024 1.857±0.013 SI 19.5.93 1200 3.35 14.8±0.4 4.950±0.022 2.283±0.027 1.861±0.012 Sl 19.5.93 1430 3.35 14.9±0.5 4.899±0.022 2.295±0.032 1.865±0.014 Cl 31.5.93 200 3.6 18.8±0.4 4.627±0.018 1.574±0.009 1.847±0.013 Cl 31.5.93 600 3.4 18.7±0.5 4.670±0.018 1.824±0.015 1.851±0.012 Cl 31.5.93 953 3.3 15.1±1.2 4. 704±0.020 2.419±0.023 1.855±0.013 Cl 31.5.93 1334 3.2 14.3±0.4 4. 719±0.02 l 2.491±0.030 1.860±0.012 Cl 31.5.93 1500 3.2 14.3±1.2 4.697±0.022 2.502±0.034 1.864±0.014 Cl 31.5.93 1620 3.1 13.8±0.5 4. 716±0.023 2.444±0.035 1.844±0.012 C2 8.6.93 200 3.6 19.4±0.5 4.623±0.018 1.537±0.009 1.848±0.017 C2 8.6.93 600 3.4 17.6±0.5 4.704±0.019 1.912±0.016 1.858±0.012 C2 8.6.93 1000 3.3 15.1±0.4 4. 763±0.020 2.306±0.023 1.852±0.017 C2 8.6.93 1200 3.2 13.7±0.4 4. 779±0.021 2.505±0.028 1.858±0.012 C2 8.6.93 1348 3.2 14.0±0.5 4.694±0.021 2.519±0.031 1.864±0.015 C3 8.6.93 20 3.7 19.6±0.5 4.599±0.017 1.483±0.009 1.848±0.012 C3 8.6.93 200 3.6 19.0±0.5 4.723±0.018 1.523±0.009 1.877±0.012 C3 8.6.93 600 3 .5 17.6±0.5 4.686±0.018 1.968±0.015 1.856±0.013 C3 8.6.93 1000 3.3 16.3±1.0 4. 725±0.020 2.397±0.023 1.857±0.013 C3 8.6.93 1200 3.2 13.8±0.4 4. 729±0.021 2.512±0.027 1.857±0.011 C3 8.6.93 1376 3.2 13.9±0.4 4. 705±0.022 2.528±0.031 1.860±0.013 C4 8.6.93 200 3.6 19.3±0.5 4.654±0.018 1.571±0.009 1.846±0.013 C4 8.6.93 600 3.4 18.4±1.2 4.683±0.018 1.979±0.016 1.844±0.014 C4 8.6.93 1000 3.3 15.0±1.l 4. 710±0.020 2.408±0.024 l.714±0.012 C4 8.6.93 1200 3.2 15.3±0.5 0.000±0.0l l 0.000±0.025 0.000±0.000 C4 8.6.93 1376 3.2 14.5±0.4 4.699±0.022 2.502±0.031 1.858±0.012 N2 17.6.93 20 2.9 18.9±0.5 4.648±0.018 l .423±0.007 1.880±0.012 N2 17.6.93 200 3.6 18.4±0.5 4. 700±0.018 1.574±0.009 1.855±0.012 N2 I 7.6.93 400 3.5 I8.5±0.0 4.714±0.018 l.566±0.0I2 1.844±0.013 N2 17.6.93 600 3.45 I 8.6±0.6 4.786±0.019 I. 703±0.015 1.848±0.0 I 2 N2 I7.6.93 700 3.45 I 9.5±0.6 4.758±0.0I9 1.565±0.0I 7 1.847±0.013 N2 17.6.93 800 3.425 I8.6±0.5 4.753±0.019 1.567±0.018 1.850±0.0I I N2 17.6.93 880 3.425 I 7.9±0.5 5.141±0.021 1.903±0.020 1.863±0.012 N3 I 7.6.93 20 3.3 I 9.2±0.6 4.654±0.0I 8 I .422±0.008 1.868±0.013 N3 17.6.93 200 3.6 I 9.0±1.0 4.685±0.0I 8 1.597±0.008 1.846±0.012 N3 17.6.93 400 3.5 I 8.9±0.5 4. 782±0.018 1.768±0.0I2 1.856±0.014 N3 I 7.6.93 600 3.55 I 9.4±0.5 4.942±0.019 1.885±0.015 1.846±0.012 N3 17.6.93 700 3.45 I8.8±1.2 4. 935±0.020 l.9I8±0.0I8 1.501±0.012 N3 17.6.93 803 3.45 I 8.8±0.6 5.07 I ±0.020 1.928±0.0I 9 1.861±0.012 FB I 8.6.93 293 3.5 I9.0±l.3 4.892±0.019 1.7I5±0.0ll l.85I±0.012 FB 18.6.93 390 3.5 I9.7±1.3 4.929±0.019 1.724±0.012 1.847±0.012 FB I 8.6.93 429 3.45 18.0±0.5 5.072±0.020 1.804±0.0I 3 1.854±0.012 FB I8.6.93 530 3.45 I 9.0±0.5 5.0I7±0.019 l.849±0.0I4 1.860±0.012 FB I9.6.93 740 3.45 I 8.3±0.5 5.056±0.020 l.938±0.0I8 1.869±0.012 NI I9.6.93 20 2.9 19.7±0.6 4.694±0.018 1.421±0.008 - NI I9.6.93 200 3.5 18.2±0.5 4.681±0.018 1.571±0.008 1.851±0.012 NI I 9.6.93 400 3.5 19.5±0.5 4.8I7±0.0I8 1.742±0.012 1.870±0.0I 3 NI 19.6.93 650 3.45 I8.7±0.6 5.044±0.020 1.933±0.017 1.850±0.012 Nl 19.6.93 800 3.4 I 9.4±0.5 5.074±0.020 1.922±0.019 1.841±0.012 NI I9.6.93 900 3.3 I 8.6±0.5 5.055±0.021 l.715±0.02I 1.850±0.013 NI 19.6.93 900 3.3 17.5±0.6 5.084±0.02 I l.789±0.02I 1.859±0.013 N2 I9.6.93 889 3.425 17.4±0.5 4. 983±0.020 1.925±0.021 1.846±0.012 Ksy 20.6.93 0 47 9.0±0.5 I3. l 72±0.050 O. I 88±0.00I 1.441±0.009 cs 20.6.93 1628 3. I 5 I4.4±0.5 4. 734±0.023 2.367±0.035 1.842±0.016 Kl 25.6.93 2I3 3.8 18.3±1.0 4.652±0.018 1.531±0.0IO 1.852±0.016 K4 25.6.93 774 3.28 I5.8±0.5 4.746±0.0I9 2.426±0.020 1.854±0.016 SI 26.6.93 20 4.3 I 8.6±0.5 4.590±0.017 1.508±0.008 1.842±0.012 SI 26.6.93 200 3. 7 17.5±0.5 4.63I±O.Ol8 1.596±0.009 1.857±0.012 SI 26.6.93 600 3.45 18.0±0.5 4.657±0.018 l. 733±0.015 1.858±0.012 142 pos date depth T Ja 4ee 3aet4He 211Ne [m) ["CJ [TU] no-8 cm3sTP ,-I 1 r10-61 no-7 cm3STP2-l Sl 26.6.93 600 3.45 18.0±o.8 4.657±0.029 1.727±0.017 - SI 26.6.93 1000 3.4 16.5±0.8 4.61 I ±0.030 1.508±0.023 l.850±0.011 SI 26.6. 93 1000 3.4 16.5±0.4 4.575±0.019 l.501±0.022 1.841±0.012 SI 26.6.93 1200 3.4 15.7±0.5 4.660±0.02 l 1.688±0.027 L853±o.Ol2 SI 26.6.93 1200 3.4 15.7±0.7 4.667±0.030 1.696±0.027 1.862±0.012 SI 26.6.93 1430 3.35 14.6±0.5 4. 721 ±0.022 2.296±0.032 1.861±0.012 SJ 26.6. 93 1430 3.35 14.8±0.4 4. 705±0. 022 2. 305±0. 032 1.339±0.010 S2 26.6.93 20 4.3 18.2±0.5 4.591±0.017 1.492±0.007 l.845±0.0l 2 S2 26.6.93 200 3.7 18.3±0.5 4.633±0.018 1.556±0.009 1.853±0.012 S2 26.6.93 600 3.45 16.7±0.5 4.686±0.018 l.949±0.016 l.848±0.013 S2 26.6.93 1000 3.4 14.9±0.5 4. 783±0.020 2.310±0.024 l.851±0.012 S2 26.6.93 1200 3.35 13.9±0.5 4.725±0.021 2.317±0.027 1.866±0.014 S2 26.6.93 1310 3.35 14.7±0.6 4.719±0.021 2.358±0.030 l.860±0.013 S2 26.6.93 1310 3.35 14.9±0.5 4.806±0.022 2.339±0.029 1.863±0.011 FB 12.7.93 0 36 10.6±1.3 2375.659±9.005 0. 197±0.001 5.451±0.035 FB 17,7.93 0 3.2 13.6±0.9 5.385±0.020 1.072±0.006 1.317±0.009 FB 17.7.93 510 3.2 17.9±0.5 5.210±0.020 1.786±0.013 l.849±0.012 FB 17.7.93 750 3.2 18.2±1.3 5 '002±0 '020 1.928±0.018 1.855±0.012 FB 17.7.93 825 3.4 18.6±1.3 5.016±0.020 1.953±0.020 1.859±0.012 S2 4. 11.94 20 6.7 20.3±0.8 4.550±0.028 l.393±0.008 1.794±0.011 S2 4.11.94 200 3.7 16.5±0.9 4.697±0.029 l .611±0.010 1.870±0.0 I 0 S2 4. 11.94 600 3.45 17.1±0.8 4.8JJ±0.030 1.851±0.016 1.870±0.012 S2 4. 11.94 1000 3.4 16.6±o.6 4.873±0.031 2.255±0.024 1.874±0.01 l S2 4.11.94 1200 3.35 14.5±1.1 4.897±0.032 2.318±0.028 1.872±0.014 S2 4. 11.94 1340 3.35 15.0±0.8 4.899±0.032 2.315±0.031 1.865±0.012 Cl 5.11.94 20 4.6 18.8±1.1 4.664±0.028 1.499±0.009 1.876±0.011 Cl 5.11.94 200 4 18.5±0.7 4.556±0.028 1.415±0.009 !.832±0.013 Cl 5.11.94 600 3.4 17.9±0.9 4.947±0.031 1.743±0.016 1.881±0.011 Cl 5.11.94 1000 3.3 17.2±o.7 5,033±0.032 2.186±0.024 1.883±0.011 Cl 5.11.94 1400 3.2 19.2±0.8 5.137±0.034 1.615±0.031 1.900±0.011 Cl 5. I l.94 1500 3.2 19.2±o.7 5.175±0.034 l.665±0.033 1.904±0.011 Cl 5.11.94 1620 3. I 19.2±0.l 4. 732±0.032 2.512±0.036 1.874±0.0ll C5 6.11.94 20 6.2 20.7±0.7 4.520±0.028 1.397±0.009 1.800±0.010 cs 6.l l.94 200 3.6 22.8±0.9 4.621±0.028 l.574±0.0 I 0 1.892±0.011 C5 6.11.94 600 3,4 21.6±0.4 4.760±0.030 1.839±0.016 0.284±0.002 C5 6.11.94 1000 3.25 20.5±0.l 4.377±0.028 1.674±0.023 l.847±0.011 C5 6.11.94 1400 3.2 19.4±0.3 4.676±0.031 l. 705±0.030 1.854±0.011 C5 6.11.94 1620 3. I 19.8±1.0 4.816±0.033 2. 152±0.036 1.880±0.012 C5 6.11.94 1620 3. I 17.8±0.7 4.832±0.033 2. 187±0.036 l.875±0.011 NI 7.11.94 200 4.2 18.6±0.7 4.625±0.028 1.463±0.010 1.828±0.011 NI 7.11.94 400 3.55 19.7±0.8 4.798±0.030 1.576±0.012 1.853±0.011 Nl 7.11.94 800 3.45 19.2±o.7 0.000±0.007 0.000±0.017 0.000±0.000 NI 7.11.94 897 3.45 17.8±0.7 4.985±0.031 1.916±0.021 1.862±0.010 FB 9.11.94 222 3.6 19.1±0.7 4. 756±0.029 1.570±0.010 1.849±0.011 FB 9.11.94 550 3.45 19.7±0.7 5.024±0.031 1.895±0.015 1.855±0.012 PB 9.11.94 649 3.45 19.0±o.7 5.006±0.031 1.901±0.017 1.863±0.013 FB 9. 11.94 677 3.45 18.8±0. 7 5.050±0.031 1.940±0.018 1.855±0.011 FB 9.11.94 760 3.45 19.3±0.7 5.034±0.031 l.934±o.Ol8 1.855±0.011 FB 9.11.94 770 3.5 19.5±0.8 5. 134±0.032 1.788±0.019 1.853±0.015 FB 9. 11.94 805 3.45 19.l±J.l 5.099±0.032 1.951±0.020 1.868±0.012 N3 9. 11.94 20 3.6 21.3±1.3 4.626±o.028 1.469±0.009 1.848±0.011 N3 9.11.94 200 3.75 16.5±1.0 4. 722±0.029 1.468±0.010 l.856±0.0ll N3 9.ll.94 400 3.55 19.5±0.7 4.811±0.030 1.683±0.012 l.850±0.010 N3 9. 11.94 700 3.45 21.1±0.9 5.147±0.032 l.902±0.018 1.863±0.011 N3 9.11.94 800 3.45 19.2±0.7 5.071±0.032 1.946±0.019 I.863±0.011 z 10.11.94 0 50 1.l±0.3 288.295±1.767 0.468±0.002 l.349±o.009 Cl 10.11.94 1400 3.2 13.6±0.8 4.843±0.032 2.466±0.032 1.876±0.011 Cl I0.11.94 1500 3.2 15.3±0.7 4.904±0.033 2.274±0.034 1.861±0.0ll C2 11.11.94 200 3.8 20.7±0.7 4.705±0.029 1.628±0.010 1.865±0.011 C2 11.11.94 600 3.4 18.5±0.7 4.829±0.030 2.001±0.015 1.678±0.010 C2 11.11.94 1000 3.3 15.9±0. l 4.701±0.030 2.445±0.025 J.855±0.012 C2 ll.11.94 1270 3.2 14.0±0.6 41.776±0.255 1.514±0.028 13.668±0.083 C2 11.11.94 1385 3.2 15.3±1.1 4. 704±0.03 l 2.530±0.032 I. 859±0.012 C3 11.11.94 200 3.8 20.1±0.8 4.633±0.028 l.514±0.010 1.851±0.014 C3 11.11.94 600 3.4 17.4±0.7 4. 735+0.029 1.895±0.016 l.870±0.012 143 pos date depth T 3ff 4He 3He/4He 20Ne [m] ["CJ [TU] [10·8 cm3STP .-11 110-61 r 10-7 cm3STP•·1 I C3 ll.ll.94 !000 3.3 15.3±0.6 4.747±0.030 2.290±0.024 l.863±0.013 C3 11.11.94 1200 3.2 18.4±0.7 4. 763±0.031 2.433±0.028 1.806±0.012 C3 11.11.94 1340 3.2 16.2±0.1 4. 701±0.03 l 2.478±0.030 l.856±0.0l I C4 I 1.11.94 200 4.2 21.2±0.0 4.679±0.029 l.498±0.010 1.860±0.011 C4 11.11.94 600 3.4 17.4±0.7 4.744±0.029 1.843±0.017 l.859±0.013 C4 11.11.94 1000 3.3 17.2±0.7 4.816±0.031 2.231 ±0.024 1.869±0.012 C4 I 1.11.94 1200 3.25 15.4±0.6 4.801±0.031 2.425±0.028 I .853±0.011 C4 11.11.94 1370 3.2 14.2±0.8 4. 746±0.031 2.536±0.032 l.866±0.014 K2 12.11.94 400 3.7 21.1±1.I 4.673±0.029 1.781±0.013 1.857±0.012 K2 12.11.94 400 3.7 19.1±1.I 4.689±0.029 1.791±0.013 1.856±0.012 K3 12.11.94 600 3.5 18.8±1.l 5.663±0.035 1.812±0.015 1.990±0.012 SI 12.11.94 20 3.7 20.4±1.3 4.633±0.028 1.507±0.009 l .850±0.01 I SI 12.11.94 200 3.8 19.3±0.5 4.656±0.028 1.525±0.008 1.859±0.010 SI 12.11.94 600 3.45 17.0±1.1 4. 740±0.029 1.910±0.015 l.859±0.010 SI 12.11.94 1000 3.4 19.8±2.0 4.840±0.031 2.252±0.024 1.864±0.013 SI 12.11.94 1200 3.35 16.9±0.8 4.881±0.032 2.314±0.027 1.868±0.010 SI 12.11.94 1300 3.35 15.l±O.l 4.844±0.032 2.317±0.030 1.872±0.011 SJ 12,11.94 1410 3.35 15.1±0.6 4.755±0.032 2.326±0,031 1.872±0.012 Sl 21.12.94 25 I 15.9±0.7 4. 709±0.038 1.433±0.009 1.884±0.017 Sl 2Ll2.94 250 3.5 19.7±0.9 4.665±0.038 l.499±0.011 1.858±0.017 SI 21.12.94 500 3.5 21.3±1.0 4.693±0.038 1.545±0.014 1.846±0.015 Sl 21.12.94 750 3.45 17.3±0.7 4. 700±0.039 l.856±0.019 1.859±0.016 SI 21.12.94 !000 3.4 15.0±0.7 4. 727±0.039 2.041±0.024 1.862±0.016 SI 21.12.94 1200 3.35 15.1±0.8 4. 736±0.040 1.899±0.027 1.855±0.015 SI 21.12.94 1400 3.4 15.1±0.1 4.872±0.041 2.359±0.032 1.887±0.020 SI 14.5.95 20 1.5 19.2±1.2 4. 727±0.038 1.400±0.008 1.920±0.018 SI 14.5,95 200 3.55 17.0±0.7 4.681±0.038 l.597±0.010 1.866±0.016 SI 14.5.95 600 3.5 15.8±0.7 4.801±0.039 1.846±0.017 1.860±0.016 SI 14.5.95 1000 3.4 14.3±0.7 4.696±0.039 2.268±0.024 l.918±0.016 Sl 14.5.95 1200 3.35 13.4±0.7 4.864±0.041 2.316±0.028 1.858±0.015 SI 14.5.95 1300 3.35 14.9±0.7 5.000±0.042 2.316±0.029 1.848±0.015 SI 14.5.95 1429 3.4 12.9±0.6 4.789±0.041 2.355±0.032 L861±0.015 K2 15.5.95 418 3.05 24.6±1.2 4.696±0.038 l.477±0.012 l.827±0.016 K5 15.5.95 1149 3.2 15.6±1.0 4.714±0.039 2.488±0.027 1.871±0.016 Cl 16.5.95 20 l.2 17.9±0.8 4.695±0.038 1. 377±0.008 1.891±0.016 Cl 16.5.95 200 3.6 17.5±0.7 4.712±0,038 1.594±0.010 l.869±0.015 Cl 16.5.95 600 3.4 16.5±0.1 4.738±0.039 l.848±0.016 1.857±0.015 Cl 16.5.95 1020 3.25 15.4±0.6 4.942±0.041 2.232±0.025 1.874±0.016 Cl 16.5.95 1420 3.2 14.6±0.8 4.975±0.042 2.451±0.033 l.949±0.016 Cl 16.5.95 1520 3.15 14.0±0.7 5.044±0.043 2.434±0.034 l.876±0.016 Cl 16.5.95 1602 3.1 13.9±1.l 4. 776±0.041 2.043±0.035 1.858±0,017 Cl 16.5.95 1624 3.1 14.0±1.0 4.824±0.042 2.281±0.036 1.851±0.016 C2 17.5.95 200 3.6 17.2±0.0 4.725±0.038 l.626±0.010 1.851±0.016 C2 !7.5.95 600 3.4 16.9±0.l 4.676±0.038 1.933±0.016 1.837±0.015 C2 17.5.95 716 3.3 16.4±0.1 4. 762±0.039 l.919±0.018 l.847±0.015 C2 17.5.95 1000 3.3 15.3±0.l 4.915±0.041 2.080±0.024 1.860±0.015 C2 17.5.95 1200 3.2 13.4±0. l 4.948±0.041 2.376±0.028 1.850±0.016 C2 17.5.95 1313 3.2 13.4±0.1 4.734±0.040 2.466±0.030 1.853±0.016 C3 17.5.95 20 l 17.3±0.8 4.719±0.038 1.388±0.009 1.900±0.016 C3 17.5.95 200 3.6 17.7±0.8 4.704±0.038 1.611±0.009 l.855±0.015 C3 17.5.95 600 3.4 17.2±0.7 4.802±0.039 l.652±0.015 0.000±0.001 C3 17.5.95 1000 3.3 15.4±1.7 4. 799±0.040 2.254±0.024 l.860±0.016 C3 17.5.95 1200 3.25 14.2±0.1 4.772±0.040 2.414±0.028 1.872±0.016 C3 17.5.95 1300 3.2 13.0±0.7 4. 797±0.040 2.470±0.030 l.849±0.016 C3 17.5.95 1391 3.2 13.4±0.9 4.710±0.040 2.391±0.032 l.850±0.016 C4 I 7.5.95 20 1.3 17.2±0.0 4.697±0.038 1.416±0.009 1.906±0.016 C4 17.5.95 200 3.6 16.4±0.7 4.654±0,038 l.605±0.010 1.853±0.017 C4 !7,5.95 600 3.4 16.9±0.6 4.706±0.038 1.763±0.016 l.850±0.016 C4 17.5.95 !000 3.3 15.1±0.7 4. 762±0.040 2.303±0.025 1.859±0.016 C4 17.5.95 1200 3.25 12.6±0.6 4. 772±0.040 2.491±0.029 1.862±0.016 C4 17.5.95 1300 3.2 15.5±1.2 4. 750±0.040 2.497±0.030 1.875±0.016 C4 17.5.95 1384 3.2 13.5±0.6 4.577±0.039 2.461±0.031 l.846±0.015 Kl 21.5.95 256 3.2 20.4±1.1 4. 706±0.038 l.548±0.010 l.882±0.017 cs 22.5.95 20 0.9 18.4±0.9 4.747±0,038 l.399±0.008 1.904±0.016 cs 22.5.95 200 3.6 17.2±0.8 4.695±0.038 1.608±0.010 1.849±0.015 144 po• date depth T 3H 4He 3He/4He :.lllNe [m) ["CJ [TUI uo-8 cm3STP 2·11 110-61 no·7 cm3STPi<-1 cs 22.5.95 600 3.4 17.9±0.8 4. 792±0. 039 1.619±0.016 1.844±0.015 C5 22.5.95 1000 3.3 14.3±0.7 S.023±0.042 2.068±0.024 1.849±0.016 cs 22.S.95 1400 3.2 12.4±0.7 4.969±0.042 2.424±0.032 1.837±0.015 cs 22.S.95 1500 3.2 13.2±0.9 4.976±0.042 2.453±0.034 l.8S6±0.016 cs 22.5.95 1626 3.1 14.6±1.0 4.762±0.041 1.983±0.035 1.848±0.016 NI 24.5.95 20 0.5 17.0±1.1 4. 794±0.039 1.403±0.007 1.889±0.015 NI 24.5.9S 200 3.6 16.8±0. 7 4.571±0.037 1.540±0.009 1.828±0.0lS Nl 24.5.95 400 3.5 15.6±1.0 4.840±0.039 1.653±0.012 1.842±0.016 NI 24.5.95 600 3.5 16.0±0.6 4.979±0.041 I. 793±0.015 1.863±0.015 NI 24.5.9S 700 3.4S 18.l±l.O S.041±0.041 J.831±0.017 1.841±0.015 NI 24.S.95 800 3.4S 16.2±0.7 5.032±0.041 1.858±0.020 l.870±0.016 NI 24.5.9S 890 3.4 17.3±0.7 4.982±0.041 1.791±0.022 1.850±0.016 Kot 27.5.95 0 80.5 0.0±0.0 1036S.3±83.9 0.050±0.001 3.282±0.028 SI 1.6.95 20 3 17.0±1.0 4.647±0.038 1.463±0.009 1.875±0.016 SI 1.6.95 200 3.6 16.1±0.9 4. 735±0.o38 I.S99±0.0IO 1.866±0.016 SI l.6.9S 600 3.5 15.0±1.2 4.903±0.040 J.824±0.015 1.845±0.015 Sl 1.6.95 1000 3.4 13.8±1.1 S.033±0.042 2.196±0.024 1.889±0.017 Sl 1.6.95 1200 3.4 13.4±0.9 4.93S±0.041 2.310±0.028 1.863±0.017 SI 1.6.95 1300 3.4 12.3±0.9 S.087±0.043 2.281±0.029 1.880±0.016 SI 1.6.95 1432 3.4 14.3±1.0 4.807±0.041 2.350±0.033 1.862±0.017 Curriculum Vitae Roland Hohmann 22 October 1965 Born in Thalwil, Switzerland 1972- 1978 Primary school in Au, Ziirich 1978 - 1980 Secundary school in Au, Ziirich 1980 - 1984 High school in Ziirich 1982- 1983 High school in Massillon, OH, USA 1985 - 1987 History and geography studies at the University of Ziirich 1987 - 1992 Study of environmental sciences at the Swiss Federal Institute of Technology (ETH), Ziirich 1992 Diploma in natural sciences (Dip!. Natw. ETH) 1992- 1997 Doctoral studies at the Swiss Federal Institute of Environmental Science and Technology (EA WAG), Diibendorf, and the Swiss Federal Institute of Technology (ETH), Ziirich, supervised by Prof. Dr. D. M. Imboden