High Alpine Lake Sediments As Chronicles for Regional Glacier and Climate History in the Upper Engadine, Southeastern Switzerland

Home , Engadine Line, Gelgia (river), Lake Sils, Lake Silvaplana, Lower Engadine, Maloja Pass

Diss. ETH No. 12705

High Alpine lake sediments as chronicles for regional glacier and climate history in the Upper Engadine, southeastern Switzerland

A dissertation submitted to the SWISS FEDERAL INSTITUTE OF TECHNOLOGY ZÜRICH

for the degree of DOCTOR OF NATURAL SCIENCES

Presented by

Christian Ohlendorf Diplom Geologe, Universität Göttingen born April 26. 1966 in Buchholz i.d.N., Germany

Accepted on the recommendation of:

Prof. Dr. H. Weissert, ETH Zürich,Examiner Dr. M. Sturm, EAWAG, Dübendorf, Co-examiner Dr. F. Niessen, AWI, Bremerhaven, Co-examiner

1998 Als ich in den Jugendtagen Noch ohne Grübelei, Da meint’ ich mit Behagen, Mein Denken wäre frei.

Seitdem hab’ ich die Stirne Oft auf die Hand gestützt Und fand, daß im Gehirne Ein harter Knoten sitzt.

Mein Stolz, der wurde kleiner. Ich merkte mit Verdruß: Es kann doch unsereiner Nur denken, wie er muß.

Wilhelm Busch Contents i Contents Contents

Abstract ...... v Zusammenfassung ...... vii Reassunt ...... x

CHAPTER 1 Introduction ...... 1 1.1 ‘Climate Change’, still a hot topic ? ...... 1 1.2 The link between climate, glaciers and lake sediments ...... 3 1.3 The high Alpine archive ...... 3 1.4 Concept and goals of the study ...... 4

CHAPTER 2 Regional overview of the Upper Engadine ...... 7 2.1 Geography ...... 7 2.2 Geology ...... 10 2.2.1 ‘General’ or pre quaternary geology...... 10 2.2.2 Quaternary geology ...... 13 2.3 Previous research ...... 17 2.4 Hydrography, Hydrochemistry and Limnology of the lakes .... 19 2.5 Climate ...... 23 2.6 Vegetation ...... 24 2.7 History ...... 25

CHAPTER 3 Methods ...... 27 3.1 Seismic survey...... 27 3.2 Coring ...... 27 3.3 Sampling ...... 28 3.4 Physical Properties ...... 28 3.4.1 Susceptibility ...... 28 3.4.2 Declination...... 29 3.4.3 Wet bulk density ...... 29 3.4.4 Grain size ...... 31 3.4.5 Layer counting ...... 31 3.4.6 Accumulation rates ...... 31 3.4.7 Suspension load ...... 32 Contents ii Contents

3.5 Geochemistry ...... 32 3.5.1 Total Carbon, Nitrogen, Hydrogen and Sulphur contents ...... 32 3.5.2 Total inorganic carbon ...... 32 3.5.3 Total organic carbon ...... 33 3.5.4 Biogenic Silica ...... 33 3.5.4 Main and Trace elements ...... 33 3.6 Mineralogy with XRD ...... 33 3.7 AMS 14C-dating ...... 34 3.8 Climate, glacier, tourism and tree ring data ...... 34 3.9 Statistical tools and data analysis ...... 35 3.10 Data storage ...... 36

CHAPTER 4 Glacial varve thickness and 127 years of instrumental climate data: A comparison ...... 37 4.1. Introduction ...... 37 4.2 Results ...... 39 4.2.1 Lithology and Chronology ...... 39 4.2.2 MAR and records of human activity ...... 42 4.2.3 MAR and climate records ...... 45 4.2.4 Comparison of suspension and sediment samples ...... 47 4.3 Discussion...... 47 4.3.1 Human impact on MAR ...... 47 4.3.2 Impact of low frequency glacier size variations on MAR ...... 48 4.3.3 Impact of high frequency variations on MAR...... 49 4.4 Conclusions ...... 53

CHAPTER 5 Evaluation of sedimentary tracers for catchment weathering and erosion modes ...... 55 5.1 Introduction ...... 55 5.2 Results ...... 56 5.2.1 Mineral composition of river- and most recent lake sediments ...... 56 River sediments ...... 58 Most recent lake sediments ...... 59 5.2.2 Mineral composition of turbidites in Lej da Silvaplauna Holocene sediments ...... 59 5.2.3 Geochemistry of the most recent and Holocene lake sediments ...... 61 Most recent lake sediments ...... 61 Holocene sediments ...... 61 5.2.4 Magnetic properties of Holocene sediments in Lej da Silvaplauna ...... 65 Contents iii Contents

5.3 Discussion...... 67 5.3.1 Recent controls on sediment distribution ...... 67 5.3.2 Turbidite mineralogy: Tracer for catchment drainage conditions ? ...... 67 The effect of sorting ...... 68 How is the signature of the source region preserved ? ...... 70 The two main mineralogical signatures ...... 70 Climatic interpretation of the turbidite record ...... 71 5.3.3 Chemical properties: Tracers for the intensity of chemical weathering ? ..73 The recent lake sediments ...... 74 Possible weathering signals ...... 76 The Holocene record of Lej da Silvaplauna ...... 78 Complications in the interpretation of element concentrations ...... 80 5.3.4 Magnetic properties: Tracers for the erosion of topsoil ? ...... 81 5.4 Conclusions ...... 84

CHAPTER 6 Interannual and interdecadal variability in the Neoglacial varve record of Lej da Silvaplauna ...... 89 6.1 Introduction ...... 89 6.2 Results ...... 91 6.2.1 Geochemistry and MAR of the sediments deposited since AD 1494...... 91 The turbidite record ...... 92 6.2.2 Comparison of the MAR record with climate (proxy) data on a regional scale ...... 93 General correlation ...... 93 Correlation in the time window AD 1738-1766 ...... 98 6.2.3 Time series analysis of the Neoglacial varve record...... 98 6.3 Discussion...... 100 6.3.1 Evaluation of geochemical sediment parameters as climate proxies...... 100 Ccarb accumulation rates as glacier size proxy...... 102 6.3.2 Event stratigraphy and constraints for a AD 1738-1766 anomaly ...... 103 Event stratigraphy...... 103 The time period AD 1738 - 1766 ...... 104 6.3.3 Regional significance of the MAR climate record with special consideration of the AD ‘1738-1766 anomaly’ ...... 105 Regional significance of the 1940s anomaly ...... 105 Regional significance of the AD 1738-1766 anomaly ...... 106 MAR and lake ice break-up dates ...... 107 6.3.4 Forcing mechanisms on varve deposition and the possibility of a non-linear response of varve formation ...... 108 Non-linear response of the MAR to climatic forcing ? ...... 108 Intervals of more pronounced periodicity ...... 109 Time series analysis ...... 110 6.4 Conclusions ...... 112 Contents iv Contents

CHAPTER 7 The Late- and Post-Glacial sediment record of Lej da Champfèr...... 115 7.1 Introduction ...... 115 7.2 Results ...... 117 7.2.1 Lithology of Lej da Champfèr sediments...... 117 7.2.2 Physics, chemistry and mineralogy of the sediments ...... 121 7.2.2.1 Lej da Champfèr core PCH 94/3 (0-1159 cm) ...... 121 7.2.2.2 Magnetic susceptibility in other cores ...... 125 7.2.2.3 Unit 4 sediments of PCH 94/3 (505-1159 cm) ...... 125 Physical and chemical properties...... 125 Mineralogy ...... 129 7.2.3 Dating of Lej da Champfèr sediments ...... 133 7.3 Discussion...... 136 7.3.1 Comparison of Late Glacial and Postglacial sediments ...... 136 7.3.2 Is the unit 3/4 boundary an isochrone ? ...... 137 7.3.3 Special characteristics of the green rhythmites ...... 141 7.3.4 Median grainsize and MAR: Glacier size proxies in the Late Glacial sediments of Lej da Champfèr ? ...... 142 7.3.5 Sediment mineralogy as a key to the glacier positions during the Late Glacial ...... 147 7.3.6 The dating problem ...... 152 7.3.7 Climatic interpretation and comparison with other records...... 156 Geomorphological background ...... 157 The Egesen (YD) ELA depression ...... 157 The YD cold interval ...... 157 Green rhythmites = YD ...... 158 Younger Dryas...... 159 Glacier retreat rates at the YD/PB transition ...... 160 Allerød ...... 162 Bølling ...... 163 7.3.8 Comparison of interannual variability in different MAR records...... 164 7.4 Conclusions ...... 165

CHAPTER 8 Conclusions and Outlook ...... 169 8.1 The use of siliciclastic sediments as chronicles of climate .....169 8.2 Proposed future research ...... 171

References ...... 175 Appendices ...... 189 Acknowledgments...... 200 Curriculum Vitae ...... 203 Abstract v Abstract

Abstract

The laminated, siliciclastic sediments of four glacier fed lakes located in a high Alpine valley (Upper Engadine) in the central Alps of southeastern Switzerland record environmental changes in a high resolution mode. Physical, chemical and mineralogical sediment properties were investigated in the present study, in order to trace these changes. The study focuses on the varved Neoglacial sediments of Lej da Silvaplauna and the laminated Late Glacial sediments of Lej da Champfèr. Forcing factors on varve formation are identified by an analysis of the modern sediments and rates of glacier size change are estimated for the Late Glacial. A model that links Late Glacial sediment accumulation rates to changes in the mode of glacier movement is proposed. The annually laminated sediments of Lej da Silvaplauna, were compared with glacier monitoring data and instrumental climate data from AD 1864 to 1990. Long- term and short-term responses to climatic change as well as anthropogenic influence can be traced separately in the varve succession. Economic development in the lake’s catchment has resulted in higher autochthonous production in the last 30 years. Autochthonous components contribute around 10% to the total amount of sediment accumulated annually since 1960, but their contribution is negligible before this date. Decadal-scale trends in annual mass accumulation rates (MAR) correlate with glacier size-variations. A stepwise, running multiple regression analysis demonstrates that interannual changes in MAR are strongly correlated with changes in mean summer temperatures, but cannot be sufficiently explained without considering summer precipitation and the number of days with snow per year. There is a ~50% probability that a given fluctuation in MAR is caused by a summer temperature anomaly. The wide range of observed correlation coefficients reveals the sensitivity of the archive to temporal variability of the climatic forcing factors. Together with a possibly non-linear response of the MAR signal to climate forcing this makes the development of transfer functions ambiguous. In a qualitative way a comparison with Swiss and European climate archives demonstrated that in most cases MAR anomalies reflect meteorological extreme conditions that are detectable on a regional scale and perhaps linked to large scale atmospheric circulation of western Europe. To estimate the influence of rainfall on sediment accumulation, for selected samples the concentration of main and trace elements, magnetic properties and the mineral composition was determined. None of these parameters allows quantitative estimates about the contribution of rainfall derived material to MAR on an annual basis. It is suggested that for high resolution studies magnetic tracers offer the highest potential if they can be calibrated with other tracers. The possibility of an early hu- Abstract vi Abstract man impact on sediment accumulation should be investigated further. XRD-analyses of river sediments and turbidites in the Holocene sequence of Lej da Silvaplauna indicate that mineralogical tracers can be employed to identify sediment sources in the geologicaly diverse catchment. In all 4 Upper Engadine lakes a drastic lithologic change marks the cessation of glacial meltwater input. Evidence is presented, that this did not happen simultaneously in all lakes. Probably Lej da S. Murezzan received meltwaters 500 years longer then two of the upstream lakes. In 6.5 m of continuously laminated Late Glacial sediments from Lej da Champfèr physical, chemical and mineralogical sediment properties were analyzed. A preliminary, floating chronology is proposed for the sediments that is based on the counting of layer couplets. Reduced sediment median grainsize in accordance with existing models is viewed as a tracer for an increased catchment glaciation. The tentative Late Glacial MAR variations disagree with the existing model, which indicates a different way of glacier movement for the larger Late Glacial glaciers. A model is proposed that ascribes low MAR that occur during maximum catchment glaciation to reduced basal shear stress (reduced erosion capacity) of cold glaciers, which is indicative of a cold, arid continental climate. Such conditions are proposed for the first 450 and the last 150 years of the YD cold interval. Sediment mineralogy was used to trace the positions and rates of movement of Late Glacial catchment glaciers. Mineralogical changes in Lej da Champfèr sediments apparently to a large extent are influenced by the glaciers in the southeastern catchment that meet contrasting lithologies on an advance. With reservation to changes in the chronology by combining grainsize and mineralogical tracers, rates of glacier size change were calculated. Accordingly it is suggested that glaciers retreated very rapidly from their YD maximum positions within 30-50 years at a rate of 70-120 m/ yr which is 5-10 times faster than for the retreat from LIA positions (13 m/yr). More- over, all tracers indicate that the YD was not a uniform cold interval. Regarding interannual variability the comparison with recent sediments reveals that the second half of the YD is punctuated by several events of unusually high MAR. This might reflect a more instable climate with larger oscillations in temperature and/or precipitation. Abstract vii Zusammenfassung

Zusammenfassung

In den laminierten, siliziklastischen Sedimenten von vier, in einem zentralalpinen Hochtal (Oberengadin) der südöstlichen schweizer Alpen gelegenen und von glazialen Schmelzwässern gespiesenen Seen, sind Umweltveränderungen mit hoher zeitlicher Auflösung aufgezeichnet. In dieser Arbeit wurden physikalische, chemische und mineralogische Sedimenteigenschaften untersucht, um diese Veränderungen nachzuzeichnen. Gegenstand dieser Arbeit sind die gewarvten neoglazialen Sedimente des Lej da Silvaplauna, sowie die laminierten spätglazialen Sedimente des Lej da Champfèr. Duch eine Analyse der rezenten Sedimente werden Faktoren, die die Warvenbildung beeinflussen, ermittelt. Die Geschwindigkeit von Veränderungen der Gletschergrössen wird für das Spätglazial abgeschätzt. Die jahresweise geschichteten Sedimente des Lej da Silvaplauna wurden für den Zeitraum AD l964-l990 mit Daten über die Zungenlängenänderung von Gletschern, sowie mit instrumentell erhobenen Klimadaten verglichen. Sowohl langfristige als auch kurzfristige Reaktionen auf Klimaveränderungen und anthropogene Einflüsse können in der Warvenabfolge getrennt voneinander verfolgt werden. Die wirtschaftliche Entwicklung im Einzugsgebiet des Lej da Silvaplauna hat in den letzten 30 Jahren zu einer erhöhten autochthonen Produktion geführt. Autochthone Sedimentbestandteile haben seit l960 einen Anteil von etwa l0% an der Gesamtmenge des pro Jahr akkumulierten Sedimentes. Vor l960 ist ihr Anteil vernachlässigbar. Auf der Zeitskala von Jahrzehnten korrelieren die Sediment-Akkumulationsraten (MAR) mit Änderungen der Gletschergrösse. Eine schrittweise durchgeführte, gleitende, multiple Regressions-Analyse zeigt, dass die jährlichen Schwankungen der MAR zwar stark mit Schwankungen der mittleren Sommertemperatur korreliert sind. Jedoch können MAR-Schwankungen nicht ausreichend erklärt werden, wenn nicht Sommerniederschlag und die Anzahl der Tage pro Jahr mit Schneefall mit berücksichtigt werden. Die Wahrscheinlichkeit, dass eine MAR-Schwankung durch eine Anomalie der mittleren Sommertemperatur verursacht wurde, beträgt ca. 50%. Die grosse Schwankungsbreite der ermittelten Korrelationskoeffizienten zeigt, dass das Sedimentarchiv sehr sensibel auf zeitliche Veränderungen der Einflussgrössen reagiert. Diese Tatsache, zusammen mit einer möglicherweise nicht linearen Reaktion des MAR-Signals auf Klimaveränderungen, schliesst die Entwicklung von Transferfunktionen aus. Ein qualitativer Vergleich des Sedimentarchives mit schweizer und europäischen Klimaarchiven zeigt, dass MAR-Anomalien in den meisten Fällen meteorologische Extrembedingungen widerspiegeln, die ein regionales Ausmass hatten und deshalb wahrscheinlich mit Veränderungen der atmosphärischen Zirkulation über Westeuropa zusammenhängen. Zusammenfassung viii Abstract

Um den Einfluss des Niederschlags auf die Sedimentakkumulation (MAR) abzuschätzen, wurde an ausgewählten Proben die Konzentration von Haupt- und Spurenelementen, magnetische Eigenschaften und die Mineralzusammensetzung bestimmt. Keiner dieser Parameter erlaubt es, quantitative Angaben über den Beitrag von durch Niederschläge erodierten Material zur MAR mit jahresweiser Auflösung zu machen. Für hochauflösende Studien scheinen magnetische Eigenschaften am geeignetsten zu sein, wenn sie mit anderen Parametern kalibriert werden können. Die Möglichkeit eines frühen anthropogenen Einflusses auf die Sedimentakkumulation sollte eingehender untersucht werden. Die röntgendiffraktometrische Analyse von Flussedimenten und Turbiditen in der holozänen Sedimentabfolge vom Lej da Silvaplauna zeigt, dass die Mineralzusammensetzung dazu benutzt werden kann, Sedimentquellen in dem geologisch sehr diversen Einzugsgebiet der Seen, zu identifizieren. In allen 4 Oberengadiner Seen markiert ein drastischer Lithologiewechsel das Ende des Eintrags von glazialen Schmelzwässern. Es werden Indizien vorgestellt, die darauf hindeuten, dass dieses Ereignis nicht in allen Seen gleichzeitg stattfand. Möglicherweise wurde der Lej da S. Murezzan 500 Jahre länger von Schmelzwässern beeinflusst als zwei der stromaufwärts gelegenen Seen. In 6,5 m durchgehend laminierten, spätglazialen Sedimenten aus dem Lej da Champfèr wurden physikalische, chemische und mineralogische Sedimenteigenschaften untersucht. Basierend auf der Zählung von Warven wird eine vorläufige Chronologie der Sedimente vorgeschlagen. In Anlehnung an existierende Modelle, werden reduzierte Korngrössen-median Werte als Phasen ausgedehnterer Vergletscherung des Einzugsgebietes interpretiert. Die vorläufigen spätglazialen MAR-Veränderungen stimmen nicht mit den existierenden Modellen überein. Dies ist ein Hinweis darauf, dass sich die grösseren spätglazialen Gletscher anders bewegten als die heutigen. Ein Modell, dass die kleinen MAR während Phasen ausgedehnter Vergletscherung mit einer Reduktion der basalen Schubspannung (reduziertes Erosionsvermögen) in Verbindung bringt wird präsentiert. Reduzierte basale Schubspannungen werden bei kalten Gletscher in kalt-aridem, kontinentalen Klima beobachtet. Aufgrund des MAR befundes werden solche Bedingungen für die ersten 450 und die letzten 150 Jahre der Jüngeren Dryas angenommen. Die Mineralzusammensetzung der Sedimente wurde benutzt, um Positionen und Veränderungsraten der spätglazialen Gletscher zu rekonstruieren. In den spätglazialen Sedimenten des Lej da Champfèr wird die Mineralzusammensetzung stark durch die Gletscher im südöstlichen Einzugsgebiet beeinflusst, die bei einem Vorstoss sehr unterschiedliche Lithologien erodieren. Vorbehaltlich Veränderungen in der Chronologie wurden Grössenveränderungen der Gletscher aus Veränderungen Abstract ix Zusammenfassung der Korngrösse und der Mineralogie der Sedimente abgeleitet. Hieraus ergibt sich, dass die Gletscher sich sehr schnell, innerhalb von 30 bis 50 Jahren, mit einer Rate von 70 -l20 m/a von ihrem Maximalstand, den sie in der Jüngeren Dryas erreichten, zurückzogen. Dies ist 5-10 mal schneller als für den Rückzug vom Maximalstand der kleinen Eiszeit (13 m/a). Alle gemessenen Parameter weisen darauf hin, dass die Jüngere Dryas keine homogene Kaltphase war. Wenn spätglaziale interannuelle MAR-Veränderungen mit denen rezenter Sedimente verglichern werden fällt auf, dass die zweite Hälfte der Jüngern Dryas durch einige Ereignisse ungewöhnlich hoher MAR gekennzeichnet ist. Dies könnte auf ein instabiles Klima mit grösseren Temperatur- und/oder Niederschlagsschwankungen hindeuten. Reassunt x Abstract

Reassunt

Ils sedimaints laminats siliciclastics da quatter lais da l’Engiandin’Ota, tuots spisgiantats d’auas da vadret, documenteschan müdamaints da l’ambiaint da cuorta dürada. In quista lavur sun gnüds stübgiads ils sedimaints, ed inpustüt sedimaints cun varvas neoglacialas dal Lej da Silvaplauna e sedimaints laminats dal temp glacial tardiv dal Lej da Champfèr cun metodas fisicalas, chemicalas, e mineralogicas. Quistas metodas permettan da documentar variaziuns aint ils sedimaints, e da stübgiar ils facturs chi influenzeschan la fuormaziun da varvas. Inplü vegna inprovà da quantifichar la sveltezza cun la quala ils vadrets han müdà lur grandezza dürant il temp glacial tardiv. La laminaziun annuala aint ils sedimaints dal Lej da Silvaplauna es gnüda congualada cun il müdamaint da la lunghezza dals vadrets e cun datas da clima masüradas pels ons 1964 fin1990. Aint illas varvas as poja disferenzchar reacziuns sün müdamaints da clima e influenzas antropologicas da differenta dürada. La prosperaziun economica aint illa regiun d’inchasch dal Lej da Silvaplauna ha manà aint ils ultims 30 ons ad ün augmaint da la producziun da material autochton. La part da material autochton han agiunt daspö il 1960 il 10% dal sedimaint accumulà düront tuot l’on. Avant l’on 1960 es quista part negligibla. Sur decennis dad ons correlescha la rata d’ accumulaziun da sedimaints (RAS) cul müdamaint da la grondezza dals vadrets. Üna analisa da regressiun multipla, fatta pass per pass, muossa cha las variaziuns annualas da la RAS correleschan fermamaing cun variaziuns da la temperatura mezdana da stà. Però i nu poulan gnir spiegadas las variaziuns in möd chi cuntantescha, sainza resguardar la plövgia da stà e’l nomber da dis düront l’on chi ha naivü. La probabiltà ch’üna variaziun da la RAS es la causa d’ üna anomalia da la temperatura mezdana da stà es intuorn da var 50%. La gronda variaziun dals coefficients da correlaziun muossan, cha l’archiv da sedimaints es fich sensitiv sün variaziuns temporalas dals facturs ils plü importants. Quist fat, insembel cun la reacziun probabelmaing na lineara dal signal da la RAS sün müdamaints da clima, nu permetta da douvrar funcziuns da transfer. Ün cungual qualitativ da l’archiv da sedimaints cun archivs da clima Svizzers e Europeans muossa, cha las bleras anomalias da la RAS represchaintan condiziuns metrologicas extremas chi han gnü üna dimensiun regionala e stan perque in connex cun müdamaints da la circulaziun aint ill’atmosfera sur l’Europa occidentala. Per pudair valütar l’influenza da la precipitaziun sün l’accumulaziun da sedimaints sun gnüdas masüradas las concentraziuns dals elemaints principals e’ls elemaints da stizis, il caracter magnetic e la composiziun mineralogica da muostras selecziunadas da sedimaint. Ningün da quists parameters nu permetta da quantifichar Abstract xi Reassunt l’erosiun chi pertocca a la RAS chaschunà tras la precipitaziun düront l’on. Per stüdis fich precis para d’esser il caracter magnetic il plü adat, però be scha las datas poulan gnir calibradas cun oter parameters. Las pussibiltats d’ün’influenza antopogena antecipada sün l’accumulaziun da sedimaints stuvess gnir stübgia plü precis. L’ analisa da diffractometria cun razs da röntgen dals sedimaints da flüm e da turbidits da las vettas dal Holocen dal Lej da Silvaplauna muossa cha la composiziun mineralogica po gnir dovrada per identifichar la funtana dals sedimaints our da la regiun d’inchasch dal lai, üna zona chi es our dal punct da vista geologic fich cumplichada. In tuot ils 4 lais da l’Engiadin’Ota documentescha ün müdamaint drastic da la composiziun dals sedimaints la finischun da l’inchasch d’aua da vadret. I vegnan preschantadas indicaziuns, chi muossan cha quists evenimaints nu chattan lö aint in mincha lai aint il listess mumaint. Probabelmaing ha survgni il Lej da S. Murezzan per var 500 ons plü lönch aua da vadret sco ils lais situats plü insü aint illa val. Aint in ün minz continuond da 6.5 m da sedimaints laminats dal temp glacial tardiv dal Lej da Champfèr sun gnü fats stüdis da’l caracter fisical, chemical e mineralogic dals sedimaints. Sülla basa da la dombraziun da varvas vegn proponü üna cronologia preliminara dals sedimaints. In vista a models existents es la reducziun da la grandezza mediana da partikels da sedimaint interpretada sco fasas da grond’extensiun da vadrets aint illa regiun d’inchasch. Ils müdamaints preliminaris dal temp glacial tardiv da la RAS corresponda cun models existaints. Quist fat muossa cha’ls vadrets gronds dal temp glacial tardiv as schmuvantaivan in different möd da quels d’hoz. Ün model chi cumbinescha las RAS plü pitschnas düront temps da grond’extensiun da vadrets cun üna reducziun da la tensiun basala (main erosiun dal vadret) vegn preschantà. Üna reducziun da la tensiun basala chatta lö pro vadrets fraids aint in ün clima fraid-arid continental. A man da las valurs da la RAS vegnan proponüdas condiziuns sco talas per ils prüms 450 e’ls ultims 150 ons dal temp da la Dryas Giuvna. La composiziun mineralogica dals sedimaints vegn dovrada per reconstruir posiziuns e ratas da müdamaints dals vadrets dal temp galcial tardiv. Aint ils sedimaints dal temp galcial tardiv dal Lej da Champfèr vegn influenzada la composiziun mineralogica fermamaing dals vadrets aint illa regiun d’inchasch meridiunal, chi erodescha pro ün avanzamaint dal vadret fich differentas litologias. Cun l’agüd da la cronologia e dal müdamaint da la grandezza dals partikels e da la mineralogia dals sedimaints sun gnüdas interpretadas variaziuns da la grandezza da vadrets. Ils resultats muossan cha’ls vadrets as han retrats fich svelt, infra 30 fin 50 ons, cun üna rata da 70 fin 120 meters l’on da lur stadi maximal chi vaivan ragiunt düront il temp da la Dryas Giuvna. Quist es var 5 fin 10 jadas plü svelt sco la retratta Reassunt xii Abstract dal stadi maximal düront l’ultim temp pitschen da glatsch (13 meters l’on). Tuot ils parameters masürats demuossan cha’l temp da la Dryas Giuvna nu d’eira üna fasa fraida homogena. Scha müdamaints dal la RAS interanuals aint il temp galcial tardiv vegnan cungualats cun quels da sedimaints recents, schi’s vezza cha la seguonda metà dal temp da la Dryas Giuvna es caracterisada tras evenimaints insolits da fich otas RAS. Quist pudess esser ün’indicaziun sün ün clima instabil cun grondas variaziuns da temperatura e/o precipitaziun. Chapter 1 1 Introduction

CHAPTER 1

Introduction

1.1 ‘Climate Change’, still a hot topic ?

The question whether the climate will change in the future always was an important issue in human societies simply because the production of food directly de- pends on it. Since food production in the industrialized countries today is highly centralized and thus the concern of only a small number of ‘farmers’ most of the people do not feel vitally dependent on climate any more. Of course in fact they are, since 50% of the world’s population live in coastal areas with an elevation of less than 5 m a.s.l. and thus would be directly affected by a global sea level rise that would be one effect of future greenhouse warming. This study deals with a region were the consequences of climate changes are thought to be particularly severe (Beniston et al., 1997), the mountains. It is known that in mountainous regions the global climate signal is strongly amplified (Beniston et al., 1997). They thus react very sensitive on climate changes, which is critical, since mountains provide direct life support for close to 10% of the world’s population (Beniston et al., 1997). The reasons for climate changes on different timescales are diverse and not completely understood. Besides orbital forcing on longer time scales, solar variability and changes in the CO2 content of the earth’s atmosphere certainly are important factors for short term climate changes. Already Arrhenius (1896) recognized CO2 as a greenhouse gas and predicted a 5-6°C warming when the atmospheric CO2 concentration would be doubled. Since today we are facing rising CO2 levels a large community of climate modelers is concerned with the question whether this prediction is realistic or not. The result of this effort is that it is more and more realized that the climate system is very complex and therefore hard to predict. Model output thus varies considerably depending on the amount of knowledge available. The effect of this is that the opinion comes into vogue that global warming will not happen (e.g. Robinson and Robinson, 1997), is nothing to worry about or mitigation might be more expensive than coping with its consequences (IPCC, 1995b). Concerning this attitude, Schneider (1990) writes: “A lot of critics for some reason overlook the fact that uncertainty is a two-edged sword: i.e. Uncertainties of physical and biological processes that leave open the possibility that the current generation of climate mod- Introduction 2 Chapter 1 els overestimated the consequences of a future warming, with equal probability in- clude the possibility that climate changes were underestimated by these models”. Meanwhile it is well known from high resolution archive studies that climate changed drastically in the past with transitional periods well within the life span of a human being (e.g. Alley et al., 1993; Dansgaard et al., 1993; Dansgaard et al., 1989; Fischer, 1996; Goslar et al., 1993). Most likely these changes were triggered by changes in the North Atlantic circulation (e.g. Lehman and Keigwin, 1992). In view of these results the possibility of a drastic future climate change is real. It seems that any human ‘attempt’ to change the composition of the atmosphere will make the system even more unpredictable. In order to break up a carefree attitude, it is necessary to inform people about the consequences of these past climate changes in the region where they are living today. This is one justification for archive studies. Another one of course is that climate models can only be tested and thus improved when sufficient high resolution climate data with a good spatial resolution are available. This is currently not the case. Data are especially scarce at high elevation sites. The most drastic climate changes happened at the end of the last glaciation which is of course way beyond the time period that is covered by instrumental measurements (at most the last 300 years). Consequently in the different high resolution natural climate archives (e.g. sediments, icecores, treerings, corals) parameters had to be identified that can act as substitute (proxy) for meteorological data like e.g. the mean summer temperature. The most frequently used proxy is the oxygen isotope ratio (d18O-value) of a material. For instance in icecores this value is used as a temperature proxy. In most of the low altitude lakes this proxy can also be used, because the sediments contain autochthonous carbonate precipitates that reflect the d18O-value and thus the temperature, of the ambient water. In addition in these lakes it is possible to use several biological and chemical tracers. Consequently for Swit- zerland during the last 25 years most of the paleoclimatic reconstructions were done in perialpine lakes. The situation is completely different in the high Alpine realm. There most of the popular climate proxies are not applicable. Most high alpine lakes do not precipitate carbonates because of low water temperatures, low primary productivity and often low carbonate concentration. Hence the well studied isotopic tracers can not be applied. Autochthonous components (algal remains, authigenic minerals) in general are also only present in minor amounts which makes the application of biological and chemical tracers difficult. Therefore, we have to look for new proxies when we want to trace the climate history in mountainous regions. The most intuitive way to do this is to search for a link with the high Alpine glacier history which is relatively well documented in the European Alps. This approach is briefly outlined below. Chapter 1 3 Introduction

1.2 The link between climate, glaciers and lake sediments

Glaciers are the most obvious witnesses of climate change. Hence everyone who lives in or often visits glaciated mountain ranges can ‘see’ that climate is changing by retreating or advancing glaciers. The glacier history thus in fact is a history of climate. More precisely, temperature and precipitation are the most important parameters that control the mass balance of glaciers. Witnesses of glacier extensions in the past are the terminal moraines, that glaciers left behind when they retreated from a maximum position. This moraine record has been intensively studied in Switzer- land (e.g. Furrer et al., 1987) and a history of glacier advances since the last glacial maximum (LGM) could be reconstructed. Two major disadvantages of this record are that (1) it is discontinuous i.e. only glacier highstands are recorded. It remains unclear how far and at which rates glaciers retreated between two glacier highstands. (2) Moraines can not easily be dated. To overcome these problems Karlén (1976) suggested to use the sediments of lakes that receive glacial meltwaters (hereafter called proglacial lakes) as archives for fluctuations of the catchment glaciers. It was recognized that variations in the composition of sediments from such lakes can mirror changes in the size of the catchment glacier. In contrast to the moraine record these lacustrine sediment records are continuous. Moreover, the chances of dating the sediments are much higher, since they are possibly annually laminated. De Geer (1912) recognized that the sediments deposited in the Swedisch proglacial lakes during the retreat of the baltic ice sheet are characterized by a rhythmic lamination. He concluded that a layer couplet that consists of a basal silt layer and a clay top represents the glacier melting cycle of one year. De Geer (1912) introduced the term ‘varve’ for such a layer couplet. Proglacial sediments thus may be easily datable by counting of their annual layer couplets.

1.3 The high Alpine archive

In this study the term proglacial lakes is used for the Upper Engadine lakes, although with 10 km distance, they are relatively far away from the glacier terminus. The Upper Engadine is a high Alpine valley (~1800 m a.s.l.) characterized by four lakes that are aligned on a southwest - northeast oriented chain, the En river (figure 2.1). Because the most important economic factor is tourism, the region is critically depending on the future climatic development, for it will influence the ecological ‘value’ and the hazard potential. As described above the study of the sedimentary Introduction 4 Chapter 1 archives of the lakes, by providing information about past climate changes, may help when decisions about projects affecting the high Alpine environment have to be made in the future. For the Upper Engadine Lej da Silvaplauna the existence of varves was proven by Leemann and Niessen (1994b). Sediments in this lake are varved between AD 1994 and 3300 BP and prior to 9400 BP, with a non-varved section in between. In Lej da Silvaplauna Leemann and Niessen (1994b) did not only use the varves as chronological tool but also as a source of paleoclimatic information. According to their model varve thickness is controlled by changes in runoff, which in turn is a function of glacier size changes and thus of changes in climate. The varved proglacial sediments from the high altitude Upper Engadine lakes in the eastern central Swiss Alps thus are high resolution archives of the glacier history. They offer the opportunity to reconstruct deglaciation of the central Alps in great detail. Based on the results of Leemann and Niessen (1994b) the prospect of the present study is to refine the reading of the climatic information, to test if the Postglacial record established in Lej da Silvaplauna can be reproduced in other lakes and to extend the varve record into the Late Glacial.

1.4 Concept and goals of the study

The ultimate goal of this study is to achieve a better understanding of how climatic signals are encoded in laminated siliciclastic sediments in order to enable a quantitative interpretation of Lej da Champfèr and Silvaplauna glacial varves. In particular this aims to elucidate how the Late Glacial drastic climate shifts are reflected in a high Alpine archive. As a first step in chapter 2 the lakes and their catchment are characterized by a compilation of geologic, hydrologic, historical and climate data. Special emphasis is put on the catchment geology because it is the most important boundary condition for the interpretation of the lake’s sedimentary record. The methods that were applied in this study are briefly outlined in chapter 3, together with information about storage and accessibility of the raw data. In chapter 4 the most recent sediments deposited during the last 130 years are compared to instrumental climate data. Anthropogenic, glaciologic and climatic controls on sediment accumulation are evaluated. The potential of annual mass accumulation rates (MAR) as proxy in paleoclimatic reconstructions is assessed. The influence of rainfall on MAR is discussed. Chapter 1 5 Introduction

Based on this, chapter 5 aims to identify parameters that allow to discriminate between glacigenic sediment components and those derived from rainfall induced soil erosion. Mineralogical, geochemical and magnetic sediment properties are analyzed in this respect. In chapter 6 a statistical analysis of the Neoglacial Lej da Silvaplauna varve record and a comparison with instrumental climate and climate proxy data for the last 500 years is presented. A better understanding of climatic forcing on the MAR signal and its regional significance is attempted. The Late Glacial sediments of Lej da Champfèr are analyzed in chapter 7. The applicability of physical, chemical and mineralogical sediment properties presented in the previous chapters as tracers for the glacier history is evaluated. The model of Leemann and Niessen (1994b) is modified for the Late Glacial sediments. A preliminary chronology and a climatic interpretation of the Late Glacial sediments is proposed. Advantages and Disadvantages about the capacity of proglacial lacustrine sediments as a source of paleoclimatic information are balanced in chapter 8. Possible future directions of research are proposed.

Chapter 2 7 Regional Overview

CHAPTER 2

Regional overview of the Upper Engadine

2.1 Geography

The Upper Engadine (Engiadin’Ota) is a high alpine valley located in the central Alps of southeastern Switzerland in Kanton Graubünden (figure 2.1). It is surrounded by 3000 to 4000 m high mountains to the northwest and southeast. The lakes in the Upper Engadine, that are dealt with in this study can be found on sheet 268, Landeskarte der Schweiz (1:50000) between: 46°24' N / 9°42' E and 46°30' N / 9°52' E, i.e. Swiss coordinates: 141.750 / 774.000 and 153.000 / 786.000. All field-names are spelled in ‘Rhätoromanisch’, the language of the Upper Engadine, according to the topographic map 268. An exception to this is the valley name, Engiadin’Ota, that is most times referred to as ‘Upper Engadine’. The rivers feeding into the lakes are named on figure 2.6, Smaller rivers are subsequently referred to by the numbers (1-21) given there in figure 2.6. The lakes form a southwest/northeast oriented chain ranging from Lej da Segl over to Lej da Silvaplauna and Champfèr towards Lej da S. Murezzan. The Engiadin is named after the river Inn (En) that originates 3 km west of Lej da Segl between Piz Grevasalvas and Piz Lunghin and flows northeastwards through all four lakes on its way to the Danube. Piz Lunghin (2780 m a.s.l.) is a triple water divide with the En draining into the black sea, the Gelgia into the north sea and the Maira into the Adriatic. The distance from the village Maloja at the southwestern end of Lej da Segl (1797 m a.s.l.) to St. Moritz on the northern shore of Lej da S. Murezzan (1768 m a.s.l.) is 15 km with an altitude difference of 29 m. This very gentle northeastern dip (7‰) of the valley bottom favoured the formation of the lakes (figure 2.2). In contrast to this rather gentle northeastern slope, to the southwest the valley is not ‘closed’ but drops very steeply towards the Val Bregaglia. The altitude difference between Maloja (1815 m a.s.l.) and Casaccia (1458 m a.s.l.), which is only 3 km further to the southwest, is already 357 m (56‰ dip). In general the slopes on the northwestern side of

Regional Overview 8 Chapter 2 Engadine (2780m)

Piz Lunghin

Zürich

Switzerland Lej da Segl da Lej (3159m)

Piz de la Margna

Lej da Champfèr da Lej

Lej da Silvaplauna da Lej Lej da S. Murezzan S. da Lej (3386m) Chapütschin (3937m) Piz Roseg

Figure 2.1: Panorama of the Upper Engadine redrawn from Kümmerly and Frey (1991). View is to the southeast into the two glacial valleys (Val Fex and Val Fedoz). The study area is highlighted on the inset map by a black square. Chapter 2 9 Regional Overview 71m C] [ Samedan [mg/l] 'Err Granite' Sandy gravel Temp Oxygen Oxygen Temperature Clayey silt and sand Silty sand and gravel 4 6 8 10 12 0246810

0 5

10 15 20 25 30 35 40

[m] Waterdepth C] [ Celerina [mg/l] Temp Oxygen Oxygen Temperature 50 50 4 6 8 10 12 0246810

0 5

10 15 20 25 30

[m] Waterdepth Northern Upper Engadine Northern Alpine foreland St. Moritz ] 3 m Sils Silvaplana Champfer St. Moritz 6 curves 10 71 29 ∗ Hypsographic Champfèr Central high Alps Water-volume [ 30 70 0 20 40 60 80 100 120 140

0 10 20 30 40 50 60 70 80 Silvaplana [m] water-depth Gravel Silt Clayey silt and sand Silty sand and gravel Calcaerous 'Bündnerschiefer' C] ] [ mg/l [

Temp Oxygen 39 61 Sils-Maria Oxygen Temperature 4 6 8 10 12 0246810

10 20 30 40 50 60 70

[m] Waterdepth ] C] [ mg/l [

Epilimnion [%] Hypolimnion [%] Temp Oxygen Oxygen Maloja Temperature 4 6 8 10 12 0246810 0

10 20 30 40 50 60 70

Southern Alpine foreland Adjacent valley to the south [m] Volume-ratio: Waterdepth

Figure 2.2: SW-NE profile of the Upper Engadine along the En (vertical scale is 10 times exaggerated). Present day lake status is characterized by temperature- and oxygen profiles at the deepest part and the hypsographic curve for each individual lake redrawn from LIMNEX (1994). Profiles of 2 deep drillings at Segl-Maria and Samedan are shown with a short lithologic description. Climate diagrams show how the continentality increases downvally from Chiavenna in the SW towards Bever in the NE. Diagrams of Lugano, Zürich and Bernina-Hospitz are shown for comparison. Regional Overview 10 Chapter 2 the Engiadin’Ota are steeper than on the southeastern side. Three glacial side valleys in the south meet the main valley which is overdeepened by around 200 m with respect to the side valleys. The Furtschellas south of Lej da Silvaplauna is a plateau region at an elevation of 2500 m. Todays topography of the Engiadin’Ota is mainly a result of the erosive work of ice-age glaciers following major tectonic lineaments (cf. 2.2).

2.2 Geology

2.2.1 ‘General’ or pre quaternary geology

The geologic history of the Upper Engadine is very complex and was subject to numerous studies in the past. Since the geology of the lake catchments dictates the lithological composition of the lake sediments it may briefly be outlined below. In the first half of this century Cornelius (1932) and Staub (1946) drew very detailed geological maps of the Upper Engadine. As a compilation of re-mapping of that area since 1970 a new tectonic map will be published soon (Spillmann, in prep.). The tectonic map (figure 2.3) shows a set of nappes in the catchment of the lakes that were overthrusted during alpine orogeny. Primarily three major tectonic units can be distinguished in the catchment. The Margna nappe (cf. Liniger and Guntli, 1988), the Platta nappe (cf. Dietrich, 1970) and the Bernina nappe s.l. (cf. Spillmann and Büchi, 1993). Major nappes and sub-nappes are separated by Mesozoic sediments (cf. Naef, 1987). Paleogeographically the Upper Pennine Platta ophiolites, today sandwiched between Middle Pennine and Austroalpine nappes, represent relics of a transform fault system in the Mesozoic central Tethys (Liguria-Piemont-Ocean) (Weissert and Bernoulli, 1985). Austroalpine nappes to the southeast and Middle Pennine nappes to the northwest represent the southern (Adria) and the distal northern (Eurasia) continental margin, respectively. However, in the Upper Engadine the tectonic position of the Platta-Ophiolites is in-between the Austroalpine Margna and Bernina s.l. nappe (Spillmann, 1993). Around 31 my the nappe structures were unconformably cut by the granitoid Bergell intrusives (von Blanckenburg, 1992). In Late Oligocene-Early Miocene lateral displacement along the Engadine Line occurred (Schmid and Froitzheim, 1993). The catchment farthest to the southwest (SW of the line Piz Lunghin-Segl-Maria- Mortel Cotschen) is made out of the Lower Austroalpine Margna nappe (see Liniger and Guntli, 1988 for discussion of its palaeogeographic position). It is composed of basement rocks, mostly orthogneisses, and schists representing the adriatic lower Chapter 2 11 Regional Overview

Abbreviations: TRIA: Granodiorit von Triangia, SU: Surettadecke, LSZ: Lanzada-Scermendone Zone, GRE: Grevasalvas, RO: Roccabella, SZ: Samedaner Zone i. allg., SC: Schlattain, PA: Padella, CL: Clavadatsch, CHA: Chastelets, LE: Chaschauna-Leverone, CG: Corna dei Gessi, OZ: Ortler Zone, QUA: Quattervalsdecke, Engadiner Dolomiten

Figure 2.3: Tektonic map of the study area modified from Spillmann (1993a). and upper continental crust with its sediment cover. Two sub-nappes (Fora and Maloja sub-nappe) can be distinguished (Liniger and Guntli, 1988). They are composed of metasediments (Fedozserie) and metagabbros (Fedozer Gabbro) in their inner parts (lower crust) and orthogneisses (Fora and Malojagneiss) in the outer parts (upper crust). Sub-nappes are separated by a thin (max. 10m) band of mesozoic clastic and carbonaceous sediments (Fedoz-Mulde). A zone of very complex slices of basement rocks, Triassic carbonates and Jurassic to Cretaceous schists, limestones and radi- Regional Overview 12 Chapter 2 olarian cherts (Trümpy, 1980) which outcrop in the Val Fex (Fexer Schuppenzone) separates the Margna nappe from the Bernina nappe s.l. in the hanging wall. The Bernina nappe s.l. represents the Austroalpine basement and consists of several sub- nappes. It is predominantly made out of Late Variscan calc-alkaline to alkaline intrusives. Minor amounts of alkaline to tholeitic meta-effusiva and high grade polymetamorphic schists, para- and orthogneisses also occur (Spillmann and Büchi, 1993). To the north, a band of ophiolites (Serpentinite, Diabas, Gabbro and Greenschists) of the High Pennine Platta nappe overlies the Sella nappe, the lower sub-nappe of the Bernina nappe (s.l.). These rocks pinch out to the south in Val Fex within the Bernina nappe s.l. between the Sella- and Corvatsch sub-nappe. The slopes southeast of Lej da Silvaplauna, up to Furtschellas and a small band heading from Plaun da Lej, on the west side of Lej da Segl, towards Piz Lunghin are composed of these ophiolites. They are overlain in the north and east by the Lower Austroalpine Corvatsch and Julier sub-nappe. Their basal thrust can be followed from below Piz Corvatsch down to Lej Ovis-chel (500m north of Surlej) by a thin band of mesozoic sediments (mainly radiolarian chert and marble). This band reappears above Segl- Baseglia, on the other side of the Engadine Line, from were it runs up to Piz Materdell. A wedge of dolomite, calcareous breccia, radiolarian chert and Bündnerschiefer between Crap da Chüern, Grevasalvas and Blaunca also belongs to this tectonic unit. Piz Corvatsch and Piz Murtel are made out of Variscan granitoids of the Corvatsch nappe. They stretch down to the shore of Lej da Champfèr and on the other side of the Engadine Line they appear again as Grevasalvas nappe at Piz Grevasalvas. Again a thin band of mesozoic sediments separates the Corvatsch granitoids from the Bernina nappe (s.str.), which is composed of gabbrodiorite, diorite and granite (e.g. in the eastern catchment, Piz Bernina, Piz Rosatsch, Ova Cotschna and St. Moritz Bad). In the western catchment Piz Güglia, Piz Polaschin and Piz Lagrev are made of granite and diorite of the Julier nappe. The slopes of Piz Nair above St. Moritz are made of garnetiferous mica schists. Mesozoic sediments (coarse dolomite and poly- gene breccia, sandstone, dolomite and mica schist) reach from the top of Piz Nair and Piz Schlattain towards Celerina. To the north granites and gneisses of the Err nappe in the northwest and the Languard nappe in the northeast occupy the north- ernmost part of the lake catchment. The Engiadin’Ota and also the Val Bregaglia are the morphologic expression of a major crustal discontinuity, the Engadine line. This sinistral strike-slip fault (Trümpy, 1977) dissects the pile of east-dipping Pennine and Austroalpine nappes. The main movement occurred during Late Oligocene to Early Miocene with a horizontal displacement of approx. 3.0 km, but also signs of neotectonic movement have been Chapter 2 13 Regional Overview detected (Pavoni, 1977). Recent studies of the kinematics point to oblique slip and block rotation along the Engadine line. For the southeastern block a seesaw-like rotation with 2.8 km uplift at Maloja no vertical movement at Samedan and 3.2 km lowering at S-chanf is suggested (Schmid and Froitzheim, 1993).

2.2.2 Quaternary geology

The theory of the Ice-Ages was first developed in Switzerland (Venetz, 1833 cited in Penck and Brückner, 1909; Agassiz, 1840; de Charpentier, 1835; de Charpentier, 1841). Therefore the European Alps are probably the best studied region of the globe in terms of their glacier (quaternary) history. At this point I may mention that, although I am totally unpatriotic, of course already Goethe envisaged the possibility of the erratics being carried to their present-day position by formerly far more extended ice masses. He mentions that in ‘Wilhelm Meister Wanderjahre’ that was first published in 1829 (Goethe, 1829). Befitting a gentleman, de Charpentier (1841) acknowledged this early observation, whereas, to my knowledge, Agassiz (1840) never said a word about this. Anyway, the first comprehensive analysis of the alpine glacier history was published by Penck and Brückner (1909). They proposed a four-fold sequence of major glaciations (Günz, Mindel, Riss and Würm) which was adopted almost all over the globe and is still used today. However, from the deep sea oxygen and carbon isotope record it is now known that perhaps more than 100 glacial cycles occurred since the beginning of northern hemisphere glaciations in the Pliocene (e.g. Raymo et al., 1990). Since the high Alps ‘suffered’ the most under this regime of glaciers, valleys like the Upper Engadine bear testimony to this especially well. The Upper Engadine drainage system was predetermined during alpine orogeny as it follows major tectonic and sedimentary structures e.g. nappe boundaries, bed- ding, fissures and faults like the Engadine Line (cf. 2.2.1, figure 2.3). When global cooling at the end of the Tertiary led to the build up of ice-sheets in the northern hemisphere the growing alpine valley glaciers also used this drainage system. The multiple flip-flop between glacial and interglacial conditions during the Quaternary implies repeated changes in the erosion mode from fluvial to glacial and back. This in the interplay with tectonic uplift, finally lead to deep erosion and the formation of the alpine valley system, that we observe today. Accordingly, probably already during the Pliocene but certainly in the Early Pleis- tocene in the Upper Engadine the Maira and Orlegna glacier shifted their drainage system from the Upper Engadine towards the Bergell (Hantke, 1983). Thus, unlike other valley glaciers the Inn-Glacier is missing its major source area due to the erosion of the southwestern end of the valley (Staub, 1952). Therefore, the main valley Regional Overview 14 Chapter 2 glacier was fed from the smaller side valleys Bernina, Morteratsch, Rosegg, Fex and Fedoz from the south and Saluver and Bever from the north. The summit of the Inn- Glacier system was located above St. Moritz and Samedan at a maximum of 2900 m a.s.l. during the Riss Glacial and probably 80-100 m lower during the Würm highstand (LGM) (Hantke, 1983). From there the major ice stream was flowing down the Inn valley. Two smaller ice streams flowed towards the Bergell and the Puschlav with transfluences at Maloja and Bernina, respectively. Although the last glaciation was not the most severe one, it is certainly the one that is most intensively studied. This is because advancing Würmian glaciers deeply eroded alpine valleys and forelands eliminating almost all traces of previous glaciations. A picture of the situation during the Last Glacial Maximum (LGM) with coa- lescing piedmont glaciers in the northern and southern alpine forelands is given in Jäckli (1962). Variations in glacier extent during the LGM and especially the chronology and mode of deglaciation in the central Alps with respect to their forelands has been an intensively studied topic for more than hundred years (Furrer et al., 1980; Penck and Brückner, 1909; Staub, 1938; Ziegler, 1876). Fliri et al. (1970) were probably the first to find evidence for an ice free Engadine valley during the Late Würmian Glacial between 31000 and 20000 BP by radiocarbon-dating of wood fragments in the ‘Bänderton of Baumkirchen’, a limnic sequence 12 km east of Innsbruck. Mean- while there is evidence that glacier extent in the eastern Alps was restricted to tributary valleys from 31 until about 24 kyrs BP (van Husen, 1997). The data compilation of van Husen (1997) also calls for a rapid ice-build up that peaked at 21 ka (‘Maximalstand’) and only lasted 3000 to 4000 years with glacier fronts some hundreds of meters behind the moraines of the peak glacial (‘Hochstand’). The subsequent meltback apparently occurred as a rapid ice decay, rather than a retreat, and lasted only hundreds maybe one thousand years. Deglaciation of the alpine forelands is indicated by ice free conditions in Lake Zürich (Lister, 1988) and Lake Lugano (Niessen and Kelts, 1989) by 14600 BP. Dates for ice-free high alpine valleys as compiled in Schlüchter (1988) range from 12580 to 14500 BP. Studies in the western Swiss Alps even suggest ice-free Préalpes before the Bølling interstadial (13000 BP) and already very small glaciers during the Bühl/Weissband stadial around 16000 BP (Schoeneich, 1997). Although these dates need further constraints they clearly point to a contemporaneous deglaciation of alpine forelands and high alpine valleys. The following Late Glacial period was punctuated by several smaller glacier re-advances before today’s glacier positions were reached at the beginning of the Holocene. Chapter 2 15 Regional Overview

Late- and Post-Glacial terminal moraines left behind by glacier re-advances in the Eastern Alps are usually classified and dated according the chronology by Patzelt (1972); Patzelt (1977) and Zoller (1977) which is based on Heuberger (1968). For the Upper Engadine this chronology is supplemented by Maisch (1981) and Suter (1981) (cf. figure 2.4). To minimize confusion arising from the use of local terminology for the different moraines following the LGM, all names are summarized in figure 2.5

Vad. Calderas

Vad. d'Agnel

Vad. Traunter Ovas

Choma Vad. Güglia Lej da S. Murezzan Mauntschas N Chalavus

Lej Marsch

Lej da Champfèr

Vad. Lagrev

Lej da Silvaplauna

Lej da Segl

Maloja Riegel Maloja Dorf Plan Fedoz Vadret da Tschierva Vadret da Morteratsch

Vadret da Roseg

Vadret dal Tremoggia

Vad. da Vad. da Vad. da Forno Fex Fedoz

Glaciers in 1973 Lakes Pollen profiles Little Ice Age (1850) Long cores River En (main drainage) Kromer stadial (Preboreal) Egesen stadial (Younger Dryas) Other rivers Daun stadial 0 1 2 3 4 km

Figure 2.4: Positions of the catchment glaciers during the Late- and Postglacial. Map modified from Suter et al. (1982) and Maisch (1992). Existing pollenprofiles and lake cores used in this study are also indicated. Regional Overview 16 Chapter 2 and integrated into the commonly used chronological framework of uncalibrated 14C years and chronozones after Mangerud et al. (1974). See also Burga (1988) and

C years BP calibrated Epoch uncalibrated (Firbas, 1954) 14 Climatostrati- Pollenzones (Mangerud, 1974) (Suter, 1981) Chronozones (Patzelt, 1977) GlacieradvancesEastern Alps Calendar years BP graphical terms Glacier advances (Stuiver & Reimer, 1993) Acheological zones Upper Engadine 0 0 Modern X Modern Sub- Middle- 1000 928 Atlan- age Late ticum Centuries AD 2000 1938 IX Roman Ironage Centruies BC 3000 3187 Bronze- Sub- age Löbben 4000 4436 VIII Boreal Rotmoos 2 Neo- 5000 5731 lithic Rotmoos 1 VII Post Glacial Holocene 6000 6821 Middle Atlan- Frosnitz VI 7000 7787 tikum

8000 8784 V Meso- Boreal lithic Venediger 9000 9979 Pre- Ova d'Err

Early IV Boreal Schlaten Beverin 10000 11179 Younger III Dryas Egesen Alp Val 11000 12917 II Allerød Pontresina Daun Bever 12000 13992 Samedan I Bølling b/c 13000 15437 Chinuos- Oldest chel Late Glacial Gschnitz I a Dryas 14000 16792

15000 17916 Palaeo- lithic Pleistocene 16000 18876 Late Pleistocene

17000 20105

18000 21484 Glacial Lascaux- 19000 Interstadial 20000 Figure 2.5: Stratigraphies used in different fields of research (geology, climatology, palynology, archeology and glaciology) shown with respect to the 14C-timesale. Calendar years in the second column were derived by calibration of radiocarbon years according to Stuiver and Reimer (1993). Chapter 2 17 Regional Overview

Furrer et al. (1987) for a comprehensive summary of Swiss glacier, vegetation and climate history and a detailed bibliography. For the Upper Engadine the Late Glacial spatial and temporal evolution of the glaciers is relatively well studied (Maisch, 1992; Suter, 1981; Suter et al., 1982). Figure 2.4, taken from (Suter et al., 1982) summarizes the current state of knowledge about the late Glacial moraine chronology. Correla- tions between different valleys in the eastern Alps are achieved by calculation of equilibrium line altitude (ELA) depressions compared to the 1850 reference level (Gross et al., 1976) and moraine-morphologic features. For instance Egesen moraines can be recognized by their fresh and blocky appearance because they did not suffer a solifluidal overprint later on (Maisch, 1981; Suter, 1981). For the Upper Engadine a still intact ice network is assumed during the Chinus- chel stade with an ice level at 2300 m above Pontresina near the confluence of Morteratsch and Roseg Glacier (Suter, 1981). Transfluence via Bernina-pass and Maloja-pass with connection to Orlegna, Maira and Albigna ice still existed at that time. Evidence is scarce for the next glacier stades Bever/Samedan and Spinas, but it suggests a decaying ice network. For the last time Bernina ice met ice from the southern side valleys (Corvatsch catchment). The peninsula Piz in Lej da Champfèr is interpreted as an ice-pushed ridge that was formed when Bernina ice flowing towards Maloja met the ice stream fed by the glaciers in Val Fex and Fedoz (Hantke, 1983). At that time the Güglia glacier terminated at Il Dschember and thus was already disconnected from the ice stream in the main valley (Suter, 1981). Following this stade glaciers retreated from the main valley and dead ice bodies were left behind on the valley floor. During subsequent glacier advances, Palüd Marscha, Alp Val, Beverin and Ova d’Err the glaciers fluctuated in the tributary valleys (Suter, 1981 and figure 2.4) and the striking alluvial fans (Isola, Baseglia, Suot l’Ova, Silvaplauna, Crap da Sass and Champfèr) that today divide the formerly continuous Upper Engadine waterbody into four lakes, were built (Hantke, 1983). All Postglacial glacier advances following the Ova d’Err stade, that was attained in the early Preboral, never reached this position again. The maximum Holocene glacier extent is marked by the moraines of the 1850 highstand.

2.3 Previous research

The unique beauty of the Upper Engadine attracted people at all times. Especially the abrupt ending of the valley at Maloja was a topic of speculation already in the Regional Overview 18 Chapter 2 early days of geologic research when it struck Arnold Escher v.d. Linth and Berhard Studer (Escher von der Linth and Studer, 1839) in the first half of the 19th century and Albert Heim later on (Heim, 1879a). Various studies from different fields of research yielded a wealth of information about different aspects of this region. Far from being complete the following compilation is a selection with limnogeological bias. For a hydrogeological survey in 1980 three deep drillings (figure 2.2) were made at Zuoz (92m), Samedan (81m) and Segl-Maria (130m) (Hantke, 1983). Unfortunately no cores were drilled so that there is neither a detailed sediment description nor are samples available. The 180 m deep drill-hole in Segl-Maria reached the bedrock (calcareous ‘Bündnerschiefer’) at a depth of 130 m (Suter, 1991, figure 2.2). In 1989 on the southwestern end of the Isola delta an exploration hole for ground-water prospection was drilled (Gilly, 1992 cited in: Gehrig, 1996). The drilling went through mostly fine-grained sediments (silts and sands) with some gravel layers as well as soil horizons and stopped at 30m depth because the ground-water production was low (Gehrig, 1996). After a geoelectric survey in 1990 a second exploration well was drilled approximately 50m north of the village of Isola in 1992 which yielded exploitable reserves of water for a possible future use (Gehrig, 1996). In the initial phase of NFP 31 six refraction seismic profiles where recorded on the Isola delta (Gehrig, 1996). Despite the relatively poor seismic results, in conjunction with the geoelectric profiles a deltaic sediment thickness of 60m below the village of Isola is assumed. Sedi- ments thin towards the western margin of the delta where dolomitic bedrock out- crops at Piz d’Islas. In the framework of NFP 31 15 cores with a maximum length of 11 m were taken on the Isola delta between 1992 and 1994. Geochemical analysis, Pollen- and 14C-stratigraphy are discussed in detail in a diploma thesis (Gehrig, 1996) and in Burga et al. (1997). Gehrig (1996) gives an estimate of 2.2 to 5.4 mm/yr for the accumulation rate on the Isola Delta. She compares these rates to accumulation rates calculated by assuming an average catchment erosion of 0.3 to 1.0 mm/yr. Making this assumption she calculates that between 7 and 52% of the material transported by the Aua da Fedoz is deposited on the delta mostly during flood events. In the same project geomorphology and glacier history of Val Fex and Val Fedoz was analyzed (Maggetti, 1994; Vogel, 1995). Geomorphologic information, is also available from several Diploma thesis (e.g. Dietrich, 1971; Seger, 1971; Wanner, 1971). A history of studies dealing with the hydrology, hydrochemistry and limnology of the Upper Engadine lakes is compiled in LIMNEX (1994). For instance Lütschg- Loetscher (1954) did a very detailed work about the ice conditions of all 4 lakes. Proceeding in the same direction recently Livingstone (1997) found that the timing Chapter 2 19 Regional Overview of ice break-up for the Upper Engadine lakes is strongly related to local and regional surface air temperatures centred on the middle of April and integrated over 4-8 weeks. Since the 1970ies several studies on sediment sequences of peat-bogs near Maloja and St. Moritz were conducted that allowed the reconstruction of the vegetation history (e.g. Heitz et al., 1982a; Kleiber, 1974; Suter et al., 1982; Zoller et al., 1996). Apart from this, concerning the lake sediments there is only one study by Züllig (1982) that traces the most recent eutrophication history of Lej da S. Murezzan. Long sediment sequences were obtained not earlier than 1990, which were then analyzed in two studies. One from Leemann (1993) dealing mainly with the physical properties of Lej da Silvaplauna Holocene sediments and another one from Ariztegui (1993) about Lej da S. Murezzan sediments with a focus on organic geochemistry. Another study that just started focuses on establishing a pollen stratigraphy for Lej da S. Murezzan, Champfèr and Silvaplauna Late and Postglacial sediments (P. Hochuli, E. Gobet, pers. comm.).

2.4 Hydrography, Hydrochemistry and Limnology of the lakes

Contrasting to most perialpine lakes, hydrochemical, physical and limnological data on the Upper Engadine lakes are scarce, probably because the high Alps are still viewed as pristine environments. This was true at best until the beginning of this century when touristic development started to affect the lake water quality especially of Lej da S. Murezzan and later on also of Lej da Champfèr, Silvaplauna and Segl. The anthropogenic induced increase in nutrient input caused a change from oligotrophy to eutrophy in Lej da S. Murezzan around 1920. This change is evident from the sediments deposited after 1900 (Züllig, 1982; Ariztegui, 1993) and was also already recognized in hydrobiological studies carried out at that time by Borner (1922) and Schmassmann (1920). Despite this discovery it took almost 50 years until the next two intensive studies (Bosli-Pavoni, 1971; Märki, 1966) confirmed the eutrophic status of Lej da S. Murezzan at the end of the 1960s. In the same studies Lej da Silvaplauna and Segl are classified as oligotrophic based on oxygen, phosphate, total phosphorus, nitrate and total nitrogen contents. However, the sediment record indicates seasonal bottom water anoxia due to increased productivity already at the beginning of the 1960s (Ohlendorf et al., 1997). Starting at the beginning of the 1960s villages in the Upper Engadine were gradually connected to sewage plants. By 1983 all sewage, except from Val Fex, was diverted from the lakes (BUS, 1985). Regional Overview 20 Chapter 2

In a comprehensive study conducted for the Kanton of Graubünden by the company Limnex the limnological situation of the Upper Engadine lakes was evaluated (LIMNEX, 1994). A compilation of the results from all previously conducted limnological studies can also be found in LIMNEX (1994). The following summary is based on this report. Morphometric, hydrologic and biologic data for all four lakes are assembled in table 2.1. The hypsographic curves and lake water temperature and oxygen profiles representing the current status of the lakes in addition are integrated in figure 2.2. All coring positions are marked in figure 2.6.

Table 2.1: Compilation of hydrographic, chemical, physical and biological data for all 4 Upper Engadine lakes.

Lej da Lej da Lej da Lej da Unit Reference Segl Silvaplauna Champfèr S. Murezzan Swiss [E] 776550 780800 781850 784900 coordinates [N] 143900 146900 149300 152150 Hydrography Altitude [m a.s.l.] 1797 1791 1791 1768 LIMNEX length [m] 4,8 3,1 1,5 1,7 Bosli-Pavoni width [m] 2,3 1,4 0,5 0,6 Bosli-Pavoni Surface area [km2] 4,1 2,7 0,5 0,78 LIMNEX Volume [106m3] 137 126,5 8,5 20 LIMNEX Maximum depth [m] 71 77 33 44 LIMNEX Average depth [m] 33,5 47 17 25 LIMNEX Catchment area [km2] 45,7 129 138 172 LIMNEX Residence time [yrs] 2,08 0,86 0,055 0,11 LIMNEX Discharge (annual av.:1974-90) [m3/s] 2,09 4,66 4,94 5,82 LIMNEX Discharge (annual sum 1974-90) [106 m3] 66 147 156 184 LIMNEX Maximum lake level fluctuation [m] 1,5 1,9 1,9 1,4 LIMNEX Catchment glaciation [%] Water chemistry und physics T max. surface waters [°C] 14,9 12,5 12,8 Bosli-Pavoni T min. surface waters [°C] 0 0 0 Bosli-Pavoni T max. bottom waters [°C] 5,5 4,6 5,2 Bosli-Pavoni T min. bottom waters [°C] 3,6 3,9 3,4 Bosli-Pavoni Transmission (green=max.) [m] 13,5 13,5 9,5 Bosli-Pavoni Min. conductivity [µS/cm] 87 95 108 Bosli-Pavoni Max. conductivity [µS/cm] 104 122 185 Bosli-Pavoni Max. oxygen [mg/l] 10 (12 in Metalimn.) 10 10 Bosli-Pavoni Min. oxygen [mg/l] 6,74 5,56 0,12 Bosli-Pavoni Karbonathärte [°fr. H.] 3.85 to 3.20 4.50 to 4.10 4.60 to 3.90 Bosli-Pavoni Posphorus (phosphate) [µg/l] 5 5 35 Bosli-Pavoni Posphorus (total) [µg/l] 11 13 43 Bosli-Pavoni Nitrate (epilimnion) [mg/l] 0,08 0,12 0,22 Bosli-Pavoni Nitrogen (total, winter) [mg/l] 0,22 0,28 0,33 Bosli-Pavoni Nitrogen (total, bottom waters) [mg/l] 0.23 to 0.33 0.28 to 0.41 0.55 to 0.67 Bosli-Pavoni Biological data Bluealgea none none none Bosli-Pavoni Mougeotia no no yes Bosli-Pavoni Biggest part of biomass Crypto- u. Chrysophyceen Diatoms Diatoms Bosli-Pavoni Biomass (total) [mg/m2] 2050 bis 35630 - 15465 to 42870 Bosli-Pavoni C ass. [mg/m2*day] 7.1 to 647 - 439 to 830 Bosli-Pavoni Biomass 0-20m [mg/l] 0.06 to 1.85 0.06 to 1.81 0.04 to 1.97 Bosli-Pavoni

All four Upper Engadine lakes belong to the dimictic, holomictic lake type (Hutchinson and Löffler, 1956). In all four lakes the first short (1-3 weeks) period of turnover usually occurs middle of May shortly after ice break up. The second, most times longer, holomixis occurs at the beginning of November. A stable summer strati- Chapter 2 21 Regional Overview

fication exists from June until October with a thermocline at 20-30m waterdepth. Inverse stratification is observed under the winter icecover from December until

Ova da Suvretta

Zürich Champfèr Inn

94/1,2,3 Engadine CH 1 93/2

Ova dal Valhun 94/4 CH 2 Silvaplauna 93/1

Surlej 21 N

92/1 92/2 94/2,3 11 92/4 90/2 . 92/6 90/1 SP 2 90/7 20 SP 1 90/4 93/5 90/3 94/1 12 Ova da Surl 93/4 19 Ova da l'Alp 18 93/1 13 ej

17 93/2+3 Ova dal Crot Sela 16 Ova da la Rabqiusa 14 15 Inn Ova da la Roda Segl-Baseglia

94/5 92/2 91/5 92/1 Lavatera 10 Fedacla 92/3 91/4 92/4 Ova dal Mulin 92/5 cores A. Leemann (1990) 9 91/3 92/6 . 93/5 Segl-Maria 8 91/2 short cores C. Ohlendorf (1993,94 94/1,2 piston-cores C. Ohlendorf (1994)

91/1 1 SEACAT-profiles (1994)

93/1 samples of river bottom 93/4 sediments C. Ohlendorf 93/2 Isola 7 Aua da Fedox 93/3 2 Inn 94/3,4 3 4 6 5 0 1.0 2.0 km Maloja

Figure 2.6: Location of all cores, SEACAT-probe profiles and river sediment samples taken in lakes Sils, Silvaplana and Champfèr during this study and the study of Leemann (1993) (see legend for explanation).

May. Due to very short spring homothermia, turnover sometimes is incomplete especially in the two deep lakes of Silvaplauna and Segl. Deep water oxygen replen- ishment during such years therefore is poor. With respect to the oxygen content of hypolimnic waters at the end of summer stratification, Lej da Champfèr and S.

Murezzan (0 mgO2/l) are classified as eutrophic, whereas Segl and Silvaplauna still have sufficient oxygen (6-7 mgO2/l) in the bottom waters and therefore are classified as oligotrophic (LIMNEX, 1994). High resolution measurements of physical water- properties done with an electric probe (SEACAT SBE 19, cf. Fischer, 1996) in this study at the end of Mai 1994 indicate a hypolimnic oxygen depletion in the northern Regional Overview 22 Chapter 2

Light transmission [%] Light transmission [%] Light transmission [%] Light transmission [%] 84 86 88 90 84 85 86 87 83.0 83.5 84.0 84.5 83 84 85 86 87 0 0 0 0 Cooler and Thermocline more turbid water from Ova da Surlej 10 Thermocline

20 10 Thermocline 10

30 Water depth [m]

40 Water depth [m] Water depth [m] Water depth [m] 20 Thermocline 20 50 10

60 Begin- ning bottom 70 30 water 30 oxygen depletion 7.0 8.0 9.0 8910 8.8 9.0 9.2 678910 Oxygen [mg/l] Oxygen [mg/l] Oxygen [mg/l] Oxygen [mg/l] SP 1 SP 2 CH 2 CH 1

Water depth Temp. [°C] Temp. [°C] Temp. [°C] Temp. [°C] [m] 4 567 567 6.4 6.8 7.1 5678 0 10

20 20 Thermocline 30 S 30 Lej da 32 measured in may '94 40 N 50 Silvaplauna 60 Lej da Champfèr 70 75

Figure 2.7: High resolution profiles of temperature, oxygen content and light transmission in the water column of lakes Silvaplana and Champfèr. Parameters were measured with an electric probe (SEACAT SBE 19) in May 1994. Positions of the profiles are also indicated in figure 2.6. Temperature profiles reveal an inclined thermocline in lake Silvaplana. Turbid water of the Ova da Surlej is recognized in the light transmission and temperature profiles. Profile CH 1 from lake Champfèr shows beginning bottom water oxygen depletion. basin of Lej da Champfèr (CH 1) already at that time (figures 2.6 and 2.7), whereas the southern basin (CH 2) does not show this feature. In contrast, other parameters like the phosphorus content and the epilimnetic productivity (chlorophyll a concentration) rather indicate oligotrophic to mesotrophic conditions for all four lakes. Phos- phorus re-dissolution is observed from anoxic sediments of Lej da Champfèr and S. Murezzan. A striking feature is the unusually thick trophogenic (10 - 20 m) zone reflecting the high transparency of the lake waters. This zone however, did not yet develop at the end of May 1994 when the SEACAT-measurements were conducted. The profiles of light transmission (figure 2.7) reveal relatively turbid water at that time, probably due to small particles still remaining in suspension from the snow- melt that mainly occurred in the middle of May. The profiles also show reduced light transmission in the upper 2 m of the water column at station SP 2. This probably reflects the influence of the Ova da Surlej (figure 2.6) that feeds water with a reduced solute load (lower conductivity) but higher suspension load (lower light transmission) and slightly lower temperature onto the lake water. Chapter 2 23 Regional Overview

On average the lakes are covered with ice for 5 months per year from the middle of December until the middle of May (Livingstone, 1997). The order of freezing and thawing is a reflection of the lakes heat budget, which in turn is controlled by hydrography (cf. hypsographic curve in figure 2.2) and wind exposure. Accordingly the shallowest and most protected Lej da S. Murezzan usually freezes first followed by Lej da Segl. The deepest and most wind exposed Lej da Silvaplauna on average is the last one to freeze. The ice break-up in spring usually happens in the order S. Murezzan, Silvaplauna and Segl. With the record ranging from 1832 until today for Lej da S. Murezzan an exceptionally long dataset of freezing and thawing dates exists. An analysis of this record by Ariztegui et al. (1992) and Livingstone (1997) revealed a long term trend towards earlier ice break-up dates. The analysis of Livingstone (1997) demonstrated that regional climatic trends and the climatic effect of large volcanic eruptions are reflected in the ice break-up dates of Lej da S. Murezzan.

2.5 Climate

The Alps are located at the southern rim of climate zone Cf (warm-temperate, humid climate) according to the Köppen-Geiger-classification (Schönwiese, 1994). There, the zone of summer-dry climate (Cs) to the south meets the temperate climate zone in the north. In the same scheme the high Alps are put into climate zone Df (humid, winter-cold climate). But as a high, central Alpine valley the Upper Engadine is characterized by a more continental summer-dry climate. This maybe best illus- trated by the large annual amplitude of mean monthly temperatures which is 19.4 °C for Segl-Maria (figures 2.2 and 2.6) compared to e.g. 18.2°C for Zürich (569m) or 17.7 for Viscosoprano (1087m) (Schüepp, 1967). The climate diagrams for the profile Casaccia-Samedan compiled in figure 2.2 illustrate, how the continentality increases down valley. Little cloud cover and intensive insolation give rise to hot summers whereas accumulation of cold air due to temperature inversions is the reason for cold winters. The prevailing westerly and northwesterly flow over Switzerland that carries moist air masses towards the northern slope of the Alps, has only little effect on the precipitation regime in the Upper Engadine. Moist air masses reach the valley mostly with a southerly flow over the pass of Maloja (Gensler, 1978). Consequently the annual sum of precipitation in Segl-Maria (1029 mm, av: 1901-60) is relatively small compared to other stations at similar altitude (Uttinger, 1966). On average maximum precipitation occurs in August, whereas a minimum is observed in January Regional Overview 24 Chapter 2

(figure 2.2). However, catastrophic rainfall events like in July and August 1987 seem to occur when humid, warm southern air masses meet very cold air masses brought to the alpine realm by a series of frontal systems from an Icelandic low. When this situation remains localized above the central eastern Alps for several days, persistent rainfall may cause catastrophic floods (Wasserwirtschaft, 1991). A peculiarity of the Upper Engadine is the ‘Malojawind’, a strong valley wind with ‘wrong’ (downvalley) direction. During sunny days in the summer month a local low pressure cell develops above the middle Engadine that starts to suck up colder air masses from the adjacent Bergell valley, via Maloja, in the morning (Urfer et al., 1979).

2.6 Vegetation

According to Walter (1990) the Alps are an interzonal mountain chain that separates a zone with temperate climate and deciduous forest (Zonobiome VI) to the north from a zone with summer dry climate and evergreen sclerophyllous woodland (Zonobiome IV) to the south. Zonation of the vegetation therefore is different in the northern, southern and central alpine realm. The inner alpine, Upper Engadine shows a pennine type (i.e. continental) forest succession (Walter, 1990). As a result of this location on the borderline of two Zonobiomes the Upper Engadine flora is amongst the most diverse ones of the Alps. It contains northern boreal elements that migrated towards the Alps from Scandinavia during the ice ages (PTT, 1960) as well as Mediterranean elements (Schmid, 1978). Due to the prevailing dry continental climate the most abundant forest tree is the larch (Larix decidua) that grows mostly on the sunny western slopes of the valley. On the shaded eastern slopes the pine (Pinus cembra) is more abundant than the larch and alpine roses are found in the undergrowth. This distribution is to a great extent conditioned by human influences especially alpine farming. Because in contrast to the larch, cembra-pine prevents the growth of grass and herbs it was cleared on sunny, easy accessible southerly slopes, giving rise to the development of larch meadows (Holtmeier, 1967b). Forest clearings started already in the early bronze ages or probably even earlier (Kleiber, 1974; Zoller et al., 1996). As a result of this, today’s timberline is situated around 2300m. However, the climatically possible timberline would be at least as high as 2500m (Holtmeier, 1967a). Following the cessation of alpine farming at the beginning of this century cembra-pine, as the more competi- tive species, starts to resettle the abandoned areas (Holtmeier, 1967b). Chapter 2 25 Regional Overview

A peculiar local climate feature controls the distribution of spruce (Picea abies). It is only found along a small zone up to 2000m between the village Segl and Lej da Staz on the northwest exposed right side of the valley within the normal, alpine rose rich larch-cembra-pine forest. That is exactly the area where a long lasting fog belt, the ‘Malojaschlange’, develops when the Upper Engadine comes under the rule of a low pressure system in the Mediterranean (Holtmeier, 1966). These small band of more moisty conditions seems to favour the growth of spruce that normally would not grow under the prevailing dry continental high alpine climate conditions. Trees that are restricted to extreme environments are Pinus mugo to swamps and Pinus montana and Alnus viridis to slopes that are endangered by avalanches, rockfalls and torrential washes and to high altitudes above the timberline (Holtmeier, 1967b). A large variety of shrubs and herbs can be found in different locations up to an altitude of more than 3000m. The flora of the Upper Engadine lakes is represented by numerous species of emerged and submerged aquatic plants. For instance in Lej da S. Murezzan (Borner, 1922) found Carex and Equisetum species as well as Chara, Nitella and Potamogeton.

2.7 History

As a consequence of its high altitude it was believed until the beginning of this century that the Upper Engadine was inhabited by humans not earlier than roman times. New evidence emerging in the course of this century falsifies this assumption. For instance neolithic influence in the Upper Engadine is inferred from several silex-blades (Conrad, 1981). The first cultivation of cereals, as evidenced by pollen analysis, was detected around 4100 BP, whereas the earliest anthropogenic influence is seen in the development of Larix meadows as early as 5600 BP (Zoller et al., 1996). So far evidence for lake dwellings that are a very common feature in almost all the lakes of the Swiss plateau during the neolithic and bronze age (e.g. Gross and Ritzmann, 1990; Staub, 1864) has never been found in high Alpine lakes. To my knowledge Ziegler (1876) is the only one who mentioned an observation made by two local people who claimed to have seen three groups of piles (14-16 or 33 piles per group) in 2.5 to 6.5 m waterdepth through the lake ice near the northeastern shore of Lej da Segl. First direct evidence of human activity in the catchment of the Upper Engadine lakes comes from knife blades found in St. Moritz and Stampa, Lej da Segl. Those findings are dated to early bronze age. In the middle bronze age settlements developed and already in the late bronze age extended settlements must have existed as is deduced from the usage of the St. Moritz mineral springs. Two barrels Regional Overview 26 Chapter 2 made of larch wood that served as water collectors at the springs were found in 1907 and later on radiocarbon dated to 3200 BP. From then on human activity is documented by numerous findings of ceramics until the first written documents appear in the late iron age (approx. 400 BC). In the year 15 BC the Engadine was assaulted by the romans (Drusus and Tiberius) and belonged to the roman empire as the prov- ince of ‘Rätia’ until 493 AD. The alpine passages of Julier, Septimer and Maloja as well as four roman altars (‘Votivaltäre’) found in Segl-Baseglia are witnesses of that time. All Upper Engadine villages were first mentioned in a document written in 1139 AD (Mair, 1995). Until 1803, the Upper Engadine came - and sometimes suffered - under the authority of different regimes. For instance during a war in 1499 all Upper Engadine villages except Segl-Maria (Zuan, 1984) burned down. From the historical record it emerges that the Upper Engadine got hit by catastrophic floods at least once every century since 1500 AD (Röthlisberger, 1991). The most disastrous flood probably occurred in 1566 when Segl, Silvaplauna, Surlej, Champfèr and the mineral springs of St. Moritz were completely destroyed (‘deeply buried under gravel’, Robbi, 1913). In 1803 the formerly independent ‘Graubünden’ joined the Swiss federal union, whereas not earlier then 1938 the language of Graubünden, ‘Rätoromanisch’ got of- ficially accepted as the fourth Swiss native language. In the 1870’s a first ‘hotel building rush’ marks the starting point of a rapid touristic development and economic growth especially for the village of St. Moritz. Well known characters like Giovanni Segantini and Friedrich Nietzsche drew their inspiration from the spectacular landscape at that time. The beginning of World War I was a first major drawback in the economic development. After a renewed improvement in the 1920’s with the Olympic winter games in St. Moritz in 1928 a second severe recession started in the 1930’s and lasted until the end of World War II. From then on, starting with the Olympic winter games of 1948, the Upper Engadine experienced a period of economic growth. Along with this, the local government tries to pursue an ecological policy attempting to balance between the tourism ‘industry’ and protection of nature (e.g.: a law against hydro- electric power generation in Lej da Segl, diversion of sewage from the lakes). Chapter 3 27 Methods

CHAPTER 3

Methods

3.1 Seismic survey

In September 1994 in total 15 km of high resolution seismic reflection profiles were recorded with a 3.5 kHz ORE transmitter in Lej das Segl, Silvaplauna and Champfèr (figure 7.1). The transmitter generates electromagnetic pulses with a maximum of 20 kW output that penetrate up to 40 m of unconsolidated, gas-free sediment with a vertical resolution of 30 cm (Giovanoli et al., 1984). The received signals were concurrently plotted on an EPC recorder after filtering and amplification. The primary signal was also recorded on a digital tape. A two way travel time of 10 ms (the distance between two reference lines on the EPC-plot) corresponds to a 7.5 m thick package of sediment, assuming a P-wave velocity of 1.5 km s-1 (Niessen et al., 1993). The positions of the track lines were determined with a sextant. The ship speed above ground was approximately 3 km/h. Coring positions were chosen after analysis of the seismic profiles.

3.2 Coring

Several short cores up to 2.5 m length have been taken from Lej das Segl, Silvaplauna, Champfèr and 5 small ponds in the watershed during 4 coring campaigns between August ’93 and October ’95. Short cores were taken with a modified ETH gravity corer, equipped with a messenger system. In September 1994 in total 12 cores up to 12 m long were recovered from Lej da Segl, Silvaplauna and Champfèr using the ETH coring system (Kelts et al., 1986) which is based on a modified Kullenberg piston-corer principle. Core locations were deduced from sextant measurements taken directly after core retrieval. All coring positions are indicated on figure 2.6. After cutting the cores into 1 m long sections they were stored in a refrig- erator at 4°C until opening. Methods 28 Chapter 3

3.3 Sampling

After accomplishment of whole core measurements (susceptibility, declination, gamma ray absorption) cores were split lengthwise into two equal halves by means of an electric router. Immediately after splitting and preparation of the cutting surfaces with a sharp blade, large-sized outline photographs were taken from the not oxidized cores. The ‘reference halves’ of the cores then were allowed to oxidize over night at 4°C, whereas the ‘working halves’ were described and sampled for water content, geochemical, mineralogical and grainsize analysis. Sampling intervals were always adjusted to the observed lithology and to the purpose they were to serve. In particular sampling intervals for the different cores used in this study are as follows: core PSP 94/3-0 (0-88.5 cm) : 122 samples comprising 1 to 5 years core PCH 94/3 (0-401 cm) : 10 samples per meter core PCH 94/3 (401-505 cm) : 20 samples core PCH 94/3 (505-1160 cm) : 50 samples per meter core PSS 94/4 (0-1063 cm) : 10 samples per meter Each sample, except those from PSP 94/3-0, represents 3.42 cm3 of wet sediment, since 0.5 cm thick pie-shaped sediment slabs were taken representing a quarter of the surface area of the core liner. For PSP 94/3-0 sample discs of varying thicknesses - adapted to the annual layering - comprising the whole surface area of the core were taken in order to calculate annual sediment fluxes with a maximum accuracy. All samples were weighed immediately after sampling and after freeze drying to calculate the water content. Exposure of the reference half to air at 4°C over night produced a scanty oxidation of the sediment surface just sufficient to let laminations appear most distinctly. Me- dium-format (6x6 cm film) high-quality photographic close-ups were thus taken one day after core opening. Each photograph, picturing 10 cm of sediment, was enlarged to a color-print of 50x30 cm from which the layer counts were conducted.

3.4 Physical Properties

3.4.1 Susceptibility

Prior to core opening magnetic susceptibility was measured with a Bartington MS2 Magnetic Susceptibility system using the MS2-C core sensor. The system has Chapter 3 29 Methods been described in detail by e.g. Nowaczyk (1991). Measurements were done in 2 cm intervals in the high precision mode of the sensor with a resolution of 2*10-6 SI. Due to the detection characteristics of the sensor susceptibility is detected in a volume (shaped like a Gaussian distribution) up to approx. 6 cm to both sides of the mid- point of the sensor. On account of this sensor characteristics measurements were continued 8 cm beyond both ends of the core sections. These additional readings are spliced with the first or last 8 cm of the adjacent core section respectively to prevent a loss of data.

Frequency dependent susceptibility (cfd) was measured in the geomagnetic lab of the Geophysical Institute of ETH on 157 samples from Lej da Silvaplauna sediments (core PSP 90/2) using a Bartington MS2 meter and a MS2-B sensor. The variation of susceptibility with the frequency of the magnetizing field is detected. This gives an indication about the presence of magnetic grains which are at the single-domain/ superparamagnetic boundary. Values are expressed as F-factors (Heller et al., 1991) and were calculated as follows: χ − χ χ = LF HF • 100 (3.4.1.1) fd χ LF

χ = Frequency dependent susceptibility, F-factor [%] fd χ LF = Susceptibility measured at a frequency of 0.47 kHz [SI] χ HF = Susceptibility measured at a frequency of 4.7 kHz [SI]

3.4.2 Declination

Declination as well as intensity of the horizontal vector of the natural remanent magnetization (NRM) were measured in the magnetic lab of the Institute for Geo- physics. Prior to opening each core was mounted in a spinner that rotates at a speed of 420 min-1. Magnetization was then measured every 2 cm with a flux gate magne- tometer (Giovanoli, 1979).

3.4.3 Wet bulk density

Wet bulk density was measured at Alfred Wegener Institute in Bremerhaven. It was determined by measuring the attenuation of gamma rays in the sediments. 137Cs was used as radioactive source and intensity was measured by a scintillator-photo- multiplier detector. The core diameter was synchronously computed by a potentio- metric system. For details of the method see Gerland (1993). Measurements were Methods 30 Chapter 3 taken every 5 mm for all cores. The attenuation of gamma rays is related to the density of the radiation-exposed material by: 1 I δ = ⋅ ln( ) (3.4.2.1) −µ ⋅ d I0

δ = density of the sediment µ = specific attenuation coefficient of the sediment d = diameter of the core I = intensity of radiation with attenuation by the sediment

I0 = intensity of radiation without attenuation by sediment (through air) Since apparatus-specific attenuation coefficients of rock and sediment have not been determined, as would have been necessary, density calibration was achieved in a different way following (Niessen, 1996). All measured core sections (n=79) were weighed and the volume of sediment they contain was calculated from the inside-dimensions of the core liner (diameter: 5.9 cm, length: 1m). Mean wet density values were calculated for whole core sections.

W s (3.4.2.2) δ s = Vs

-3 δ = mean wet density of a whole section (g cm ) s = weight of a section (g) W s = inside-volume of the liner (cm3) V s 1.9 Densities obtained from equation 3.4.2.2 1.8 y = -3.1851x + 3.7256 R2 = 0.9801 should then be a function of the mean gamma

] 1.7 3 cm

/ 1.6 counts (mean scintillation detector voltage of all g [ 1.5 measurements on a core section) measured on a 1.4

Density given section. Figure 3.1 shows the regression 1.3 1.2 between gamma ray attenuation (mean detec- 1.1 tor voltage) and density. The correlation coeffi- 0.60 0.64 0.68 0.72 0.76 0.80 Mean detector voltage [V] cient of 0.98 demonstrates the validity of this Figure 3.1: Density calibration. The mean de- method for density calibration. All densities were tector voltage of 79 measured core sections computed by gamma ray attenuation is plot- calculated using the following equation, derived ted against the density computed by weigh- from linear regression: ing of whole cire sections. The equation computed by linear regression is used to calculate δ =−3.1851⋅ I + 3.7356 (3.4.2.3) density values from the high resolution gamma-ray logs with a resolution of 5mm. I = detector voltage Chapter 3 31 Methods

3.4.4 Grain size

A Laser-Particle-Analyzer, type Galai CIS-1 (Co. LOT, Darmstadt, GER), was used to determine the grain size distribution of the sediments < 150 µm. Samples for grain size analysis were taken by scratching off some tens of milligrams of sediment from the cleaned surface with a 0.5 cm wide spatula. The samples were then suspended in

15 ml of 10 % H2O2 to remove minor amounts of organic matter and to promote deflocculation. After one week reaction time and 30 seconds of ultrasonic treatment, aliquots were taken from the homogenized suspension with a Pasteur pipette and diluted in the measuring cuvette with demineralized water to a concentration appropriate for the laser system. Measured diameters were converted to sphere volumina which then were classified into 20 different categories between 0.5 and 150 µm by the system software. The software also calculates median, mean, mode and standard deviation of the grainsize distribution following the formulas given in McManus (1988). To estimate the accuracy of this method always 3 replicates were measured for each sample. The standard deviation of the 3 measurements always was below 5 %. The error introduced by inaccuracies in the sampling procedure was assessed by twofold re-sampling of 20 samples representing different lithologies. Replicate measurements of these samples also yielded a standard deviation below 5 %. Grain size classification refers to the Udden-Wentworth scheme (Udden, 1914; Wentworth, 1922).

3.4.5 Layer counting

Layer counts and thickness measurements were conducted on high quality color- prints (50x30 cm) each comprising 10 cm of sediment. A digital caliper was used to measure the thickness of individual layers. Clay layers were used to define layer boundaries.

3.4.6 Accumulation rates

Bulk mass accumulation rates (MAR) and Component accumulation rates (CAR) were calculated from layer thicknesses using the algorithm described in Niessen et al. (1992) as follows: (3.4.6.1) MAR = s ⋅ρs ⋅(1 −Φ)

w ⋅ρ Φ= w (Berner, 1971) (3.4.6.2) w ⋅ρw + (1 − w) ⋅ρs Methods 32 Chapter 3

MAR ⋅ C CAR = (3.4.6.3) 100

s = sedimentation rate (cm) between two dated points (usually = layer thickness) = dry density of the sediments (here: 2.65 g cm-3 = mostly quartz) ρs = pore-water density (1.0 g cm-3) ρw Φ = Porosity after (Berner, 1971) w = water content of the sediments (weight %) C = concentration of a given component (weight % of the total dry weight) MAR and CAR are always given in mg or g dry sediment per cm2 per year.

3.4.7 Suspension load

Suspension loads were measured using three-liter water samples that were freeze- dried and weighed after centrifugal separation. The suspension concentration is given in mg per liter.

3.5 Geochemistry

3.5.1 Total Carbon, Nitrogen, Hydrogen and Sulphur contents

Total carbon (TC), total nitrogen (TN), total hydrogen (TH) and total sulphur (TS) contents of the freeze-dried, homogenized bulk samples were measured using gas chromatography (CHNS Elemental Analyzer, Model EA 1108, Co. Carlo Erba Instru- ments). For analysis a small tin container was charged with 5 to 15 mg of sediment and measured against a reference standard (sulfanilamide) which had the following concentrations: C: 41.84 %, N: 16.27 %, H: 4.68 %, S: 18.62 %. A reference standard was measured every 10 samples.

3.5.2 Total inorganic carbon

The total inorganic carbon (TIC) content was measured using a coulometer (Cou- lometric Inc., 5011 CO2-Coulometer). For analysis 100 mg of sediment were reacted with a 2 mol/l HCl solution at 80 °C and measured against 100% Na2CO3 as a reference standard. Chapter 3 33 Methods

3.5.3 Total organic carbon

The total organic carbon content (TOC) was calculated by subtracting the total inorganic carbon portion (3.5.2) from the total carbon content (3.5.1).

3.5.4 Biogenic Silica

The biogenic silica content was measured using a leaching method where biogenic silica is extracted with 1 M NaOH at 85 °C. Dissolved silica content is concurrently monitored during the leaching process by molybdate-blue spectrophotom- etry. See Müller and Schneider (1993) for details.

3.5.4 Main and Trace elements

For 31 samples from Lej da Silvaplauna sediments, main- and trace-element concentrations were measured after total disintegration with ICP-OES at EAWAG in Dübendorf. Disintegration of 50 mg of sediment for each sample was done using an autoclave system (Heinrichs et al., 1986) following the procedure described in Heinrichs (1990); Heinrichs and Herrmann (1990); Ohlendorf (1993). The elements Na, K, Ca, Mg, Al, Sr, Fe, Mn, Zn, Cu, Cr, Ni, were measured with ICP-OES (Co.: Spectro).

3.6 Mineralogy with XRD

The mineral composition of sediment 350 samples was determined by X-ray diffraction 300 y = 152.51x + 13.406 R2 = 0.857

(XRD) of smear slides. For smear slide prepara- 250 tion approx. 3 mm3 of sediment were suspended 200 with acetone in a mortar and ground to <63 µm. Dolomite [cps] 150 RD

The suspension was then transferred to a glass X 100 sample holder were it slowly dried out. This procedure caused a preferred orientation of platy 50

0 minerals which allowed their identification, 0.00 0.20 0.40 0.60 0.80 1.00 1.20 1.40 1.60 1.80 2.00 Coulometer C inorg [%] even if only small quantities were present, by Figure 3.2: XRD intensities of the 2.89Å do- enhancing the base reflections. Diffraction pat- lomite peak plotted versus the content of car- terns were recorded with a Scintag XDS 2000 at bonate carbon computed by coulometry. The good positive correlation indicates that sedi- 45 kV and 40 mA with Cu Ka radiation between mentary carbonate in lake Champfèr mainly 2 and 70°2Q. Intensities (peak heights) of the do- is detrital dolomite. Methods 34 Chapter 3 lomite (virtually no calcite is present in the sediments) 100% peak at 30.958°2Q (d=2.886 Å) were plotted versus total inorganic carbon contents (figure 3.2). The observed positive correlation (r2=0.86) between those two parameters validates the utilization of peak intensities as semiquantitative measures for the concentration of a mineral in a given sample relative to another sample (cf. Fanning et al., 1989).

3.7 AMS 14C-dating

For AMS 14C-dating in total 17 samples were taken from the sediments of Lej da Champfèr. From these samples, 8 contained too little organic matter for dating and 9 were dateable. The dates obtained are given in table 7.1 and figures 7.3 and 7.12a. Dating was done at the Institute for Middle Energy Physics, ETH-Hönggerberg and at University of Utrecht using the AMS dating technique. Table 7.1 depicts the properties of the dated samples and the measured ages. 14C-ages are given in years BP, i.e. years before AD 1950. All dates are corrected to a d13C-value of 25‰. 14C- dates were converted to calendar-years using the program CALIB 3.0.3A (Stuiver and Reimer, 1993). Due to the scarcity of organic matter in late glacial sediments, dating of terrestrial macrofossils was not always possible (table 7.1). One date (sample 10 in table 7.1), due to the absence of macrofossils, was done on bulk organic matter after decarbonatization.

3.8 Climate, glacier, tourism and tree ring data

Climate data from the meteorological station in Segl-Maria, located at 1802 m a.s.l. (46°26‘ N; 09°46‘ E) adjacent to Lej da Silvaplauna (figure 2.6 and 4.1) were compiled from (SMA, 1864-1994). Data on the length variations of the Morteratsch and Roseg Glaciers in the two valleys east of the Upper Engadine main valley (figure 4.1), were obtained from (SANW, 1986-1994). Data on the development of tourism in the villages of Segl, Silvaplauna, Maloja and Surlej was compiled from the local tourist offices (Segl, 1993/94; Silvaplauna, 1993/94). The tree ring data sets were kindly provided by Prof. F. Schweingruber from WSL Birmensdorf. Chapter 3 35 Methods

The dates of lake ice break-up in Lej da S. Murezzan were kindly provided by Livingstone (1997). Temperature and precipitation data from long-term European instrumental climate records were taken from Jones and Bradley (1995). Seasonal and annual means of the following records were considered: Temperature records from Berlin, Stockholm, Geneva and Central England. Precipitation records from England+Wales, Hoofdorp, Milano, Padua, Lund, Marseille and Paris. The dataset for the Central England temperature series is from Manley (1974) and Parker et al. (1992). The complete datasets are also available via the Internet (ftp://ftp.ngdc.noaa.gov/paleo/ climate1500ad/).

3.9 Statistical tools and data analysis

All data smoothing was done using non weighted running arithmetic means, averaging over minimum 3 to maximum 101 data points depending on the dataset they were applied to (exact numbers are always denoted on the respective graph). In chapter 4 and 6 running correlation coefficients between two datasets were calculated applying the technique described by Dean and Anderson (1974). The correlation window is either 11 or 21 data points (indicated on the respective figure). All datasets were normalized and smoothed with a 5 point running mean prior to correlation analysis to account for uncertainties in the respective chronologies. The same correlation technique was used for the stepwise multiple regression analysis. 11 data points were used to compute the correlation coefficient for each step. All values were normalized prior to comparison applying the following equation: x −∅ z = (3.9.1) σ with: z = normalized value x = measured value ø = arithmetic mean of all data points s = standard deviation Spectral analysis was done using both, a Lomb transformation and the Blackman- Tuckey (BT) fourier transformation. For the BT power spectra a ‘Welch’ window (Welch, 1967) was used. All time series analysis, except the Lomb transformation, Methods 36 Chapter 3 were done with the program ‘AnalySeries’ (Paillard et al., 1996) provided on the Internet. The Lomb transformation was done with a program from U. Wortmann (pers. comm.).

3.10 Data storage

The rawdata files can be obtained from the author in digital form or as a printout upon request. Chapter 4 37 Comparison with climate data

CHAPTER 4

Glacial varve thickness and 127 years of instrumental climate data: A comparison

Published in: Climatic Change 36(3-4), 391-411.

4.1. Introduction

Scientists searching for records of climate in the recent past (Holocene) follow two major pathways: 1) Instrumental and historical records document climate over the last decades and centuries. The information potential of these data has long been recognized (e.g. Brückner, 1889). Lang (1885) was perhaps the first who compared temperature and precipitation data with length variations of glaciers. 2) Sedimen- tary archives contain information on climate and climatic change over decadal to millennial scales. When de Geer (1912) recognized varves as annual sedimentary layers produced by seasonal variations in glacial meltwater production, he opened the possibility of using varved proglacial lake sediments as high resolution records of past climate. Recent work has demonstrated that glaciers react very sensitively to changes in climate variables (e.g. Patzelt and Aellen, 1990; Tangborn, 1980). The need to calibrate sedimentary archives against instrumental climate data, as has been done for historical data by Pfister (1995) in order to understand present and past forcing mechanisms in the Earth’s climate system, has been emphasized by Glenn and Kelts (1991). Attempts to link biological, geochemical, sedimentological or geophysical signatures of annually laminated (varved) lake sediments to meteorological or hydrological parameters have been made by numerous authors (e.g. Granar, 1956; Østrem, 1975; Gilbert, 1975; Perkins and Sims, 1983; Renberg et al., 1984; Leonard, 1985a; Leonard, 1986; Leemann and Niessen, 1994a; Desloges, 1994; Itkonen and Salonen, 1994; Teranes and McKenzie, 1995). Apart from three of these studies, where non-glacial varves were investigated, all of the above-mentioned studies used varves from proximal proglacial or ice contact lakes for comparison with meteorological data. Using proximal archives seems problematic since Østrem (1975) discovered that there is no simple relationship between discharge (reflecting weather conditions, i.e. mainly temperature and precipitation) and suspended sediment transport (controlling sediment accumulation rates, i.e. varve thickness in a proglacial Comparison with climate data 38 Chapter 4 lake) in a proximal glacial meltwater stream. Nevertheless Østrem (1975) presumed that a simple relationship can be found farther downstream in the drainage system in distal lakes. Distal proglacial lakes therefore may provide a high-resolution record of glacial erosion and runoff and thus of palaeoclimate (Karlén, 1981; Leonard, 1985b; Leemann and Niessen, 1994b). For instance, Leemann and Niessen (1994b) demonstrated that long-term varve thickness variations during the last 3500 years in Lej da Silvaplauna reflect glacier-size variations, which are caused by northern hemisphere temperature changes (inferred from the Greenland ice core d18O-record). Furthermore, most of the above-mentioned studies were carried out on time intervals of 20 to 40 years and often rely on climate data from stations that are sometimes more than 30 km away from the lake. Due to this limitation, spatial variations in climate have to be disregarded and long-term dynamic glacier response to climatic change cannot be investigated. It is known that various glacial processes and weather conditions influence the sedimentation in a proglacial lake on different time scales, ranging from annual to millennial. Consequently these are recorded superimposed on each other in proglacial lake sediments. In many cases proglacial varve thickness correlates positively with summer temperatures (Leonard, 1985a; Leemann and Niessen, 1994a; Desloges, 1994). Other important climatic parameters are rainfall and snow cover albedo (Leemann and Niessen, 1994a). In order to obtain reliable palaeoclimatic information from proglacial lacustrine sediments it is therefore necessary to disentangle the signals recording these different processes. Accurate dating of the deposits is an additional problem that arises when varve characteristics are correlated with instrumental climate data. The controversy about the annual nature of Skilak Lake “varves” may exemplify this problem (Perkins and Sims, 1983; Stihler et al., 1992; Jerikowic and Sonett, 1993). Increased autochthonous production due to recent eutrophication may mask the climate signal that is produced only by allochthonous sediment input from the catchment glacier. To isolate the climatic signal that can be deduced from the yearly variations in varve characteristics, autochthonous and allochthonous input need to be traced separately. In this paper, we assess the potential of Lej da Silvaplauna glacial varves for reconstructing past climatic change, by comparing varve thickness variations with climate, glacier and tourism data. Lej da Silvaplauna is located in the Upper Engadine valley in south-east Switzerland (figure 4.1). We compare the varves with a 127 year long instrumental climate data series from a climate station located 1 km south-west of the lake. 112 years of glacier monitoring data from Morteratsch and Roseg Glacier (14 and 8 km east of Lej da Silvaplauna, respectively) is also used for comparison Chapter 4 39 Comparison with climate data

N Zürich

Inn Lake St. Moritz Switzerland Engadine Flax Inn Lake Champfèr

Lake Silvaplauna (1791 m a.s.l.) Ova da Roseg Glacier Inn monitoring

Meteorological station Sils-Maria (1802 m a.s.l.) Vadret da

Aua da Fedoz Tschierva Inn Lake Sils Fedacla Vadret da Morteratsch Vadret da Roseg

Locations of cores Vad. da taken in Sep. 94 Güz Vad. dal Location of core Vad. da Tremoggia PSP 94/3-0 Fedoz 01234 km Vad. da Fex

Figure 4.1: Map of the study area in the Upper Engadine (south-east Switzerland). Lakes in light grey, glaciers in dark grey. Glacier monitoring sites at Morteratsch and Roseg Glaciers and the position of the meteorological station in Sils-Maria are indicated by arrows. The black square in the inset map indicates the location of the study area.

(figure 4.1). Lej da Silvaplauna is a small (2.7 km2), deep (78 m), distal (10 km from glacier terminus), proglacial, high altitude (1791 m a.s.l.) lake. 8% of the catchment area is covered by three small glaciers (0.89, 1.33 and 2.49 km2) south of the lake (Maisch, 1992). This study area hence has the advantage of a high-resolution geological record and the availability of long instrumental climate and glacier data series from nearby stations. This chapter focuses on the results obtained from the core PSP 94/3-0 (figure 4.2) taken at a water depth of 74 m (position marked with a triangle in figure 4.1).

4.2 Results

4.2.1 Lithology and Chronology

The sediments of the 90.0 cm short core PSP 94/3-0 (figure 4.1) consist mainly of quartz, feldspars (mostly albite, only a little orthoclase), micas (mainly chlorite and Comparison with climate data 40 Chapter 4 muscovite) and minor amounts of dolomite, calcite, amphiboles and serpentine minerals. Of the four lithological units described in Leemann and Niessen (1994b), the topmost two units are present in the short core (figure 4.2).

MARtot 0 100 200 300 depth (cm)PSP 90/1unit Age BP depth (cm)PSP 94/3-0unit Age AD depth (cm)PSP 94/3-0unit Age AD 1 1994 -2 -1 1987 Year [mg cm yr ] 100 1 1987 1984 385 10 1 1974 200 1945 20 1964 300 10

925 1954 400 30 2 1015 1888 1957 1944 500 1725 40 1934 1945 600 20 2 1924 50 700 1816 2775 1914 2875 60 1808 2 800 2965 1909 1904 3835 1894 900 70 30 4620 3 1888 5845 1884 1000 7830 80 1879 8755 1700 1874 1100 9415 -2 -1 4 1864 [mg cm yr ] a 90 1679 b c 1864 d Black, organic carbon-rich, clastic varves 0 500 1500 Brown-gray clastic varves MARturbidites Black/brown/white-layered, organic carbon-rich laminites Coarse silt to sand turbidites Figure 4.2: Schematic core section of a) a piston-core taken in 1990 with 14C dates in years BP and the lithological units defined in Leemann (1994b) (redrawn after Leemann, 1994b) b) a 90 cm gravity-core taken in 1994 (this study) and c) the topmost 38.5 cm of the gravity-core shown under b) that was used for comparison with climate data. d) Total mass accumulation rates (MARtot) inferred from varve thickness measurements, smoothed with a

5-point running mean, and the mass accumulation rate of turbidites (MARTurbidites) identified in the varve section. Dates are given in years AD based on varve counts.

Unit 1 (from 0 to 16 cm sediment depth) consists of black clastic varves with an mean organic carbon content of 2.1% and an mean biogenic silica content of 10.5% (table 4.1), with abundant pennate diatoms. This unit was deposited between 1960 and 1994 based on the varve chronology. Unit 2 (from 16 down to 90 cm sediment depth) comprises brown-gray clastic varves with a mean organic carbon and biogenic silica content of 0.5% and 1.1%, respectively (table 4.1), with a few centric diatoms. This unit was deposited between 1724 and 1959 based on the varve chronology. Chapter 4 41 Comparison with climate data

Table 4.1: Mean, minimum and maximum total mass accumulation rates (MARtot) in units 1 and 2 of the varve record as well as mean organic carbon and biogenic silica content.

Lithological Year AD Mean Min. Max. Mean Mean unit MAR MAR MAR organic biogenic (yr) (yr) carbon silica content content -2 -1 -2 -1 -2 -1 [mg cm yr ] [mg cm yr ] [mg cm yr ] [%] [%] 1 1994-1960 85 35 (1970) 232 (1993) 2,1 10,5 2 1959-1864 159 36 (1914) 386 (1889) 0,5 1,1

There is a good correlation of the observed lithological succession with existing cores described by Leemann (1993). Characteristic sedimentological features such as marker beds (characteristic turbidites), eye-catching varve successions and susceptibility swings served as correlation levels. Leemann (1993) showed that the varve counting is consistent with event dating, e.g. the 1987 flood event (Bundesamt für Wasserwirtschaft, 1991) and the 1986 and 1963 137Cs peaks due to Chernobyl and atomic bomb testing, respectively. The dates Leemann (1993) obtained can therefore be transferred to the core used in this study based on the sedimentological correlation. This study focuses on the topmost 38.5 cm of core PSP 94/3-0, comprising the last 130 years, i.e. from 1864 to 1994 (figure 4.2), because instrumental meteorological data are available for this period. To obtain maximum time resolution starting at the preserved sediment/water interface this core section was sampled varve by varve (whenever possible) using the clay layers to define varve boundaries. Because a defined volume of sediment was taken as sample, mass fluxes could be measured directly in mg dry sediment per cm2 per year. MAR were also calculated from the varve thickness measurements using the algorithm described in Niessen et al. (1992), assuming a dry density for the mineral matter of 2.65 (mostly quartz). The results of the two calculation methods are in agreement (figure 4.3a). On average the MAR calculated using the varve thickness measurements is 5% higher than the directly measured MAR. This overesti- mate may be due to gas trapped in the sediment. The varve thicknesses measured in this study also can be directly compared with those from Leemann (1993) (figure 4.3b). Varve thickness, expressed as total mass 2 accumulation rate per cm per year (MARtot), shows a generally decreasing trend from 1864 to the present (figure 4.3b, table 4.1). However, the curve of Leemann (1993) does not seem to show the same decreasing trend (stippled curve in figure

4.3b). The points where the largest offsets between the two MARtot curves are observed (marked with arrows in figure 4.3b) occur where rapid increases in the water content of the sediment were detected (figure 4.3c). The effect of rapid shifts in the Comparison with climate data 42 Chapter 4

0 100 200 0 100 200 300 30 40 50 60 70 80 1994 -2 -1 -2 -1 Year [mg cm yr ] [mg cm yr ] [%] 1984

1974

1964

1954

1944

1934

1924

1914

MARtot 1904 MARtot calculated this study MARtot 1894 MARtot weighed Leemann (1993) Water content 1884

1874

1864 a b c

Figure 4.3: Checking the method. a) Comparison of total mass accumulation rate (MARtot)weighed, derived from sampling annual layers, with MARtot calculated from varve thickness measurements. b) Comparison of

MARtot from this study with MARtot from the work of Leemann (1993) calculated from varve thickness measurements. All curves are smoothed with a 5-point running mean. c) Water content of the sediments measured with annual or two-year resolution. Arrows indicate points where rapid increases are observed. See text for further explanations.

water content on MARtot values was underestimated by Leemann (1993), because high resolution measurements of the water content were not available. Considering this, the remaining offset between the two MARtot curves is due to the different coring positions in the lake.

4.2.2 MAR and records of human activity

Because no calcite is produced in Lej da Silvaplauna surface waters, diatom-frustules and organic matter represent the major portion of particles produced within the lake by biological activity. Due to increased autochthonous production during recent years in Lej da Silvaplauna, the climate signal that is created by the allochthonous sediment input (MARall) can be obscured. In order to estimate the contribution of autochthonous particles (MARaut) to the bulk mass of particles accumulated per unit area per year at the lake bottom (MARtot) the biogenic silica, total organic carbon, total nitrogen and total sulfur contents of each annual sediment layer were measured (cf. section 4.1). Subtraction of MARaut from MARtot (figure 4.4b) should then yield the climatic signal (MARall). Chapter 4 43 Comparison with climate data Year 1994 1984 1974 1964 1954 1944 1934 1924 1914 1904 1894 1884 1874 1864 00 ] 3 -1 tot all yr 0 -2 25 MAR MAR refer to MAR with refer 00 [mg cm all 2 Comparison of mass a) 0 ) 15 tot 00 and MAR 1 tot 0 5 0 45 -1` , representing all autochthonous components , representing curve. MAR -1 00 Tourist tot 3 ∗

3 yr 0 Segl-Maria Silvaplauna, Surlej, Champfèr -2 overnight stays yr 10 15 0 b 0 2 from the MAR from aut 15 aut 0 MAR 51 ) from the total mass accumulation rate (MAR the total mass accumulation ) from aut 0 0 1. ] -1 yr .5 -2 0 total sulfur All parameters given in [mg cm 0 . 0 .4 0 .3 0 .2 0 total nitrogen .1 0 0 Data on tourism since 1935 and the effect of subtracting MAR Data on tourism since 1935 and the effect 3. b) ). 0 aut 2. total organic carbon 0 1. 0 . 0 12 Subtraction of the mass accumulation rate of autochthonous components (MAR Subtraction of the mass accumulation a biogenic silica 48 16 0 Year 1994 1984 1974 1964 1954 1944 1934 1924 1914 1904 1894 1884 1874 1864 Figure 4.4: Figure accumulation rates of biogenic silica, total organic carbon, total nitrogen, total sulfur and the sum of all four in mg cm total sulfur and carbon, total nitrogen, organic accumulation rates of biogenic silica, total and without (only allochthonous material) autochthonous components, respectively. The dashed line represents the arithmetic mean. MAR curves are smoothed with a 5- the arithmetic mean. MAR curves are line represents The dashed autochthonous components, respectively. and without (only allochthonous material) point running mean. accumulated at the lake bottom (MAR Comparison with climate data 44 Chapter 4

All autochthonous components show an increase above a background level starting around the year 1960 (figure 4.4a). The shapes of the biogenic silica, total organic carbon and total nitrogen curves are almost identical. The accumulation rates rise above a background level to a first maximum between 1959 and 1975. Total organic carbon and total nitrogen values show only a slight increase (from around 0.8 and 0.1 to 1.3 and 0.2 mg cm-2 yr-1, respectively), whereas the biogenic silica-accumulation rate increases remarkably (from around 2.0 to 6.9 mg cm-2 yr-1). Between 1975 and 1986 all values are generally high even though they fluctuate considerably. Maxima are observed in 1982 and 1986. From 1986 to 1987 component accumulation rates show a pronounced decrease (from a total of 18.2 to 7.8 mg cm-2 yr-1). From then until today they increase consistently (total organic carbon and total nitrogen more strongly than biogenic silica). During the time of generally low values (1864 - 1959) distinct maxima occur in 1888, 1889 and 1921. Also, the time interval from 1942 to 1950 shows higher fluxes and higher variability (figure 4.4a). The total sulfur curve displays the most pronounced increase. Before 1959 the total sulfur accumulation per year in the sediments is virtually zero (below the detection limit). Between 1959 and 1975 it increases to a maximum of 1.0 mg cm-2 yr-1. From then onwards, the total sulfur accumulation decreases stepwise until today (zero) with peaks in 1982 and 1986 (0.8 mg cm-2 yr-1). Based on the four measured autochthonous components, three intervals can be clearly distinguished (figure 4.4a). Between 1974 and 1994 the highest MARaut rates are observed (mean = 12.7 mg cm-2 yr-1). Between 1960 and 1973 the mean rate is 7.9 mg cm-2 yr-1. From 1864 to 1959 low rates are observed with an mean value of 3.0 mg cm-2 yr-1. Accordingly, autochthonous components make only a minor contribution to MARtot. Before 1960 their contribution is less than 2% and can therefore be neglected. Only after 1960 a noteworthy contribution can be detected (between 9.3 and

15.0% of MARtot), in which biogenic silica is by far the most important constituent (78%) followed by total organic carbon (16%), total sulfur (4.5%) and total nitrogen (1.5%). To isolate the climatic signal that can be deduced from the yearly variations in MARall, the autochthonous components were subtracted out. The resulting MARall curve shows an offset of approximately 10% to lower values in the eutrophic part after 1960 (figure 4.4b). Nevertheless, the shape of the curve remains the same as it was before modification. Prior to 1960 the two curves are identical, indicating that the annual sediment production by autochthonous activity was insignificant during that time.

During the time period investigated in this study, the MARall curve exhibits a generally decreasing trend (section 3.1). However, exceptionally high MARall values oc- Chapter 4 45 Comparison with climate data

0 50 100 150 200 250 300 0 100 200 300 1994 Year [mg cm-2 yr-1] [mg cm-2 yr-1] 1984 Morteratsch Glacier 1974 Roseg Best fit 4-th order Glacier Polynomial 1964

1954 MAR = MAR - 1944 det all MARall Glaciertrend 1934

1924

1914

1904

1894

1884

1874 Glacier retreat [m] 1864 -2000 -1000 0 Figure 4.5: Retreat of Morteratsch Glacier (since 1882) and Roseg Glacier (since 1895) compared with the mass accumulation rate of allochthonous material (MARall). A best-fit curve matching the glacier-retreat curve was used to subtract the long-term glacier trend from short-term fluctuations. The curve on the right-hand side shows the detrended MAR curve that results after subtraction (MARdet). The dashed line represents the arithmetic mean. MAR curves are smoothed with a 5-point running mean. cur from 1885 to 1897 and from 1943 to 1951. From 1951 to 1989, MARall is commonly low, but starts to rise again between 1989 and 1994. Between 1897 and 1943, MARall fluctuates around the long-term mean of 139 mg cm-2 yr-1, apart from the minima that occur between 1909 and 1916. From 1864 to 1880, MARall decreases stepwise and then starts to rise towards a maximum in 1889.

4.2.3 MAR and climate records

The Morteratsch Glacier has retreated continuously from 1882 until today, but at varying rates (figure 4.5). Between 1882 and 1937, a moderate retreat rate of 11.2 m/ yr on average was observed, similar to that from 1968 until today. The highest retreat rate of the monitoring period was measured between 1938 and 1967, in which the most dramatic retreat with a mean of 39.0 m/yr, was observed between 1958 and 1967 (1938-1957: mean=25.9 m/yr). For Roseg Glacier a similar trend is observed but with an offset of 10 years. The three meteorological parameters used for comparison with MAR in this study are mean summer air temperature, cumulative summer precipitation and the num- Comparison with climate data 46 Chapter 4 ber of days with snow (figure 4.6). The mean summer air temperature (May-Sept.) for the entire observation period from 1864-1990 is 8.8 °C. The highest summer temperatures were recorded between 1942 and 1953 (mean=9.5 °C). Other temperature maxima occurred in 1868 and 1911 (11.5 °C). The coolest summers occurred in 1912 (7.4 °C) and 1984 (7.3 °C).

0 100 200 8.0 9.0 10.0 400450 500 40 50 60 70 80 90 1994 -2 -1 Year [mg cm yr ] [˚C] [mm] [d] 1984 Cumulative Mean summer summer Number of 1974 air temperature precipitation days with MARdet (Mai-Sept.) (June-Sept.) snow yr-1 1964

1954

1944

1934

1924

1914

1904

1894

1884

1874 [mg cm-2 yr-1] 1864 0 500 1500 MARTurbidites

Figure 4.6: Detrended mass accumulation rate curve (MARdet) in comparison with mean summer air temperature (May-Sept.), cumulative summer precipitation (June-Sept.) and the number of days with snow per year. The dashed lines represent the arithmetic means. All data shown are smoothed with a 5-point running mean. The MAR of all turbidites occurring in the core-section is also shown.

The mean (1864-1990) cumulative summer precipitation (June-Sept.) is 433 mm. Rainfall values show a high interannual variability, varying between 195 and 737 mm. Rainfall maxima occur in 1890, 1946, 1954, 1960 (737 mm) and 1987 (691 mm). Generally moderate values (no extremes) occur between 1906 and 1923. Minima were measured in 1886, 1949, 1959 (195 mm) 1961, 1962 and 1983. The mean number of days with snow per year is 63. The range is quite large (23 to 106) and interannual variability is also high. Periods with a generally low number of snow days are 1864-1895 and 1940-1951, with a mean of 49. Intermediate values are reported during the periods 1900-1913, 1923-1937, 1957-1975 and 1980-1990 (mean=69). Periods with an exceptionally high number of snow days are 1916-1922 and 1974-1979 (mean=81). Chapter 4 47 Comparison with climate data

4.2.4 Comparison of suspension and sediment samples

Suspension samples were taken from the Fedacla - the glacial outwash stream and main tributary entering Lej da Silvaplauna (figure 4.1) - on a warm day during the ablation season of the glacier in August 1994 (normal glacier melting) and during a heavy rainfall event in September 1994. The suspension loads were 74 and 322 mg/l, respectively. In addition, two sediment samples from core PSP 94/3-0 were taken, one from the regularly varved section (varve of the year 1900) and one from the turbidite layer caused by a catastrophic rainfall event in 1987 (Bundesamt für Wasserwirtschaft, 1991). The mineral composition of all four samples is virtually identical. They all consist of micas (mainly muscovite and chlorite), quartz, feldspars (mostly albite, only little orthoclase) and minor amounts of amphiboles.

4.3 Discussion

During the last 127 years, MAR in proglacial Lej da Silvaplauna was affected by changes in the following variables: human activity in the catchment area, the size of the catchment glacier and the meteorological conditions (i.e. the temperature during the ablation season, the precipitation during summer/fall and the amount of snow falling during accumulation period). These three factors do not vary independently of each other and their influence on MAR changes through time. In the following the different influencing factors are discussed separately.

4.3.1 Human impact on MAR

Biogenic silica, organic carbon, nitrogen and sulfur are mainly produced within the lake itself by biogenic activity in the water column and at the sediment/water interface. The most important component, biogenic silica, is certainly purely autochthonous (diatom frustules). Most of the organic carbon is presumably also autochthonous. Allochthonous organic matter accounts for less than 40% of the total organic carbon flux as estimated from the total organic carbon content of the pre- eutrophic sediments. This is also true for nitrogen, which co-varies with organic carbon. Sulfur compounds are probably formed diagenetically in anoxic pore-water. The production and consequent preservation of these autochthonous components is controlled by the trophic status of the lake, which in turn is controlled by the nutrient input. figure 4.4b shows that the measured increase in the accumulation of all autochthonous components coincides with the increase in tourism in the villages of Comparison with climate data 48 Chapter 4

Segl (Sils, 1993/94), Silvaplauna, Maloja and Surlej (Silvaplana, 1993/94). An increase in anthropogenic nutrient input as a result of increased tourism is presumably the most important cause of the recent eutrophication of the lake. Bosli-Pavoni (1971) described Lej da Silvaplauna as an oligotrophic lake at the beginning of the 1970’s on the basis of the water chemistry and phytoplankton and zooplankton assemblages. However, sediments deposited in 1970 already showed signs of anoxia (black color) and higher amounts of biogenic silica, total organic carbon and total sulfur, indicating increased productivity. The increase in sulfur accumulation is apparently caused by the better preservation potential for sulfur due to anoxic conditions within the hypolimnetic pore-waters. Sulfur is presumably fixed in a reducing environment (sediment color turns from brown-gray to black in 1960) in the sediment as iron sulfide (Jones, 1982; Sweeney and Kaplan, 1973). Hence, although recent eutrophication drastically changes the lithology (sediment color, porosity), the effect on MAR is very small. There is a clear dominance of allochthonous clastic sediment input into Lej da Silvaplauna (section 3.2), which can be clearly recognized in figure 4.4b. The MARall curve that results when the autochthonous components are subtracted out therefore represents a climatic signal.

4.3.2 Impact of low frequency glacier size variations on MAR

It has been shown by Leemann and Niessen (1994a); Leonard (1985a); Leonard

(1986), that decadal to century-scale variations in proglacial lacustrine MARall are closely related to up-valley glacier extent. Higher sediment availability (i.e. more glacial erosion during times of larger glaciers) leads to a long-term increase in MARall.

In order to identify this low-frequency glacier activity signal in the MARall curve we compared it with glacier-size variations (figure 4.5). Since no data on variations in glacier-size were available from the three glaciers in the Lej da Silvaplauna catchment (Vadret da Fex, Vadret dal Tremoggia, Vadret dal Güz), we used data from Morteratsch and Roseg Glaciers situated in adjacent valleys east of Lej da Silvaplauna (figure 4.1), for comparison purposes. This seems reasonable because regional differences in climate are fairly small at these sites, so that a similar glacier history can be assumed. We calculated a best-fit curve that matches the observed glacier-size curve as closely as possible (figure 4.5). The polynomial function obtained was used to subtract out the effect of the long-term trend in glacier size from the MARall measurements. The resulting detrended MAR curve (MARdet) (figure 4.5) is smoothed and fluctuates around a mean value of about 113 mg cm-2 yr-1. Chapter 4 49 Comparison with climate data

In Lej da Silvaplauna, MARall decreases from 1864 until today due to glacier retreat (figure 4.5). Today, the catchment glaciers produce less rockflour than, e.g., during the Little Ice Age, when glaciers were up to 60% larger than today (Hoelzle, 1994).

Consequently highest Holocene MARall values in Lej da Silvaplauna are observed during the Little Ice Age glacial maximum between 1790 and 1850 (Leemann and Niessen, 1994b). In response to the cold summers from 1909 to 1916, 75% of Swiss glaciers were advancing around 1920 (Patzelt, 1985). Adjacent to Lej da Silvaplauna, Roseg Gla- cier was advancing, whereas Morteratsch Glacier was still retreating, but at a lower rate. In the warm decade of 1940-1950 and immediately after that time, virtually all of the Alpine glaciers were retreating due to increased, temperature-induced ablation. Roseg Glacier retreated dramatically from 1949 to 1957 (Maisch, 1992) at a mean rate of 49 m/yr. In contrast, the exceptionally dramatic retreat of Morteratsch Gla- cier did not start until 1958 and lasted until 1967, with a mean retreat rate of 39 m/yr (figure 4.5). As this dramatic retreat was a response to the exceptionally warm summers of 1942 to 1953, Morteratsch Glacier can be estimated to lag air temperature changes by approximately 15 years. In contrast to this, the corresponding Roseg Gla- cier lag time is only 5 years. This is in agreement with the response time given by Gamper and Suter (1978). The response time of the 3 small glaciers (Vadret da Fex, Vadret dal Tremoggia, Vadret dal Güz) feeding their meltwaters into Lej da Silvaplauna is approximately 0 to 4 years. The response of these small glaciers to the warm summers starting in 1942 is seen in increased MARall values between 1943 and 1951. During this rapid retreat, a considerable amount of suspended material was released and glacigenic sediments became exposed and were entrained by periglacial reworking in the glacial stream. This immediate response is in agreement with results published by Nesje et al. (1995). MARall decreased drastically around 1951 and has remained low until today. The rapid decrease may be explained by sediment depletion in the source regions induced by glacier retreat. The same phenomenon is also described by Desloges (1994).

4.3.3 Impact of high frequency variations on MAR

A high-frequency variation pattern is superimposed on the low-frequency variations in MARall caused by changes in the size of the catchment glacier. The suspension load of the glacial outwash stream accounts for the majority of sediment input to Lej da Silvaplauna. All other sediment sources are of minor importance. The sediment deposited annually in the profundal zone of Lej da Silvaplauna represents the amount of fine-grained material released from the catchment in the respective year. Actually, sediment trap studies in proglacial Lake Oeschinen (Leemann and Niessen, Comparison with climate data 50 Chapter 4

1994a) revealed that an annual couplet represents the sedimentation of only the summer, since frozen catchment conditions during the winter interrupt the delivery of detrital particles to the lake, causing a hiatus. The suspension load of the glacial outwash stream is related to runoff (Leemann and Niessen, 1994a), even though the relationship between these two hydrological parameters is complex (Collins, 1990). If only yearly averages are considered and not single runoff events, a linear correlation between runoff and suspension load is observed (Østrem, 1975). In partially glaciated catchments, runoff is controlled by both rainfall and glacial melting. Con- sequently, peak runoff years leading to high sediment input into the lake (high annual MARdet) can be related to either exceptionally high precipitation or exceptionally high mean summer temperatures causing increased glacial melting, or both. This is why, for instance Leemann and Niessen (1994a) found a decreasing correlation coefficient between runoff and summer temperature in the catchment of Lake Oeschinen with increasing summer precipitation. An attempt to trace these two different sources of runoff (and suspension) in the sediments of Lej da Silvaplauna was made. The mineral composition of suspension samples taken from the glacial outwash stream under two characteristic hydrological conditions was compared with selected sediment samples representing these conditions (cf. section 3.4). The fact that all the samples showed practically the same mineral composition allows us to conclude that heavy rainfall events result in the mobilization of dominantly glacial and fluvioglacial deposits (Østrem, 1975; Lister, 1984). Therefore we cannot use mineralogical methods as a tool to distinguish glacial meltwater MARdet peaks from those caused by heavy rainfall events. Since there is no means of distinguishing these two factors we used a different approach to assess the importance of rainfall for changes in MARdet. The three most important meteorological parameters for the generation of runoff are: 1) mean summer (May - Sept.) temperature; 2) cumulative summer (June - Sept.) precipitation; and 3) the number of days with snowfall per year. These parameters were compared to the MARdet curve of Lej da Silvaplauna sediments from 1864 to 1990 (figure 4.6). In order to display how the degree of correlation between each meteorological parameter and the MARdet varies through time, the running correlation coefficient technique described by Dean and Anderson (1974) was used. The resulting continuous curves (figure 4.7a) show at what time the correlation between two variables is strongly positive (+1), inverse (-1) or not present at all (0). In the following these curves are discussed in further detail. The correlation curves (figure 4.7a) indicate that mean summer temperature is the most important factor controlling MARdet variations, as has been already established elsewhere (Desloges, 1994; Leemann and Niessen, 1994a; Leonard, 1985a). This was Chapter 4 51 Comparison with climate data 99% =0.47 interval 2 confidence 82% > r det MAR Multiple regression =0.47. 2 0.0 0.5 1.0 -1 / det ) and mean summer temperature, cumulative summer ) and mean summer temperature, det MAR Number of days with snow yr 43% < -0.5 -1 0 1 / 26% > 0.5 det MAR Cumulative summer precipitation -1 0 1 The curve shows the result of a running multiple regression analysis (see text for explanation). The 99% confidence of a running multiple regression The curve shows the result b) rrr / det 51% > 0.5 Mean summer air temperature MAR ] -1 0 1 ] -1 -1 yr yr -2 -2 det [mg cm [mg cm 1500 MAR Turbidites Curves of the running correlation coefficient (r) between detrended mass accumulation rate (MAR coefficient (r) between detrended Curves of the running correlation a) 500

ab MAR 0 100 200 0 curve is smoothed with a 5-point running mean. curve is smoothed with a 5-point running det Year 1994 1984 1974 1964 1954 1944 1934 1924 1914 1904 1894 1884 1874 1864 Figure 4.7: Figure precipitation and the number of days with snow. The numbers at the bottom of each curve denote the percentage of data points with a correlation coefficient greater than 0.5. The coefficient greater a correlation of data points with The numbers at the bottom of each curve denote the percentage with snow. and the number of days precipitation MAR interval is shown in grey. The number at the bottom of the curve depicts the percentage of data points plotting above the confidence level of r of data points plotting above the confidence The number at the bottom of the curve depicts the percentage interval is shown in grey. Comparison with climate data 52 Chapter 4 especially true for the time intervals 1874-1884, 1898-1900, 1906-1922, 1927-1930, 1939-

1957 and 1962-1970, when the correlation between MARdet and mean summer temperature is very high (close to 1.0). Remarkably, the best correlation between mean summer temperature and MARdet occurs during years with exceptionally high or low values of mean summer temperature. During these times, changes in MARdet seem to be caused entirely by changes in mean summer temperature.

When the correlation of MARdet with mean summer temperature is low, the correlation with summer precipitation is generally high. This holds true for the time periods 1885-1897, 1930-1934, 1958-1961 and 1974-1977. During these intervals, mean summer temperature remains close to its long-term mean of 8.8 °C. Apparently, the concurrent occurrence of small positive deviations from the long-term summer precipitation mean (433 mm) is sufficient to yield a positive correlation between MARdet and summer precipitation. Again, a good correlation between summer precipitation and MARdet is usually observed during years with exceptionally high summer precipitation. It appears that past extreme meteorological conditions seem to have a strong effect on MARdet (figures 6, 7a), which is in agreement with the findings of Saarnisto (1983), who reported that hydrological and meteorological extreme conditions are often reflected in unusually thick or thin varves in Lake Päijänne, Finland. An inverse correlation is observed with the number of days with snowfall. This can be explained by the albedo effect. Fresh snow has a high albedo and therefore, glacial meltwater production is reduced during years of frequent snowfall (Wendler and Weller, 1974). This in turn leads to lower sediment input into the lake and hence to lower MARdet. It has been shown that mean summer temperature is not the only factor controlling MARdet, as summer precipitation and the number of days with snow play an important role too. In particular, the influence of rainfall on runoff and therefore on

MARdet, increases with decreasing glacier size (Collins and Taylor, 1990; Fountain and Tangborn, 1985). In order to obtain an estimate of how important each of the three meteorological parameters is, for explaining the observed variability in MARdet signal, we first normalized all three parameters. Secondly, we conducted a running, stepwise multiple regression analysis (figure 4.7b). Thirdly, the mean contribution as well as the range of variability was estimated from the stepwise removal of parameters (table 4.2). By applying a one-tailed t-test, it has been confirmed that a significant correlation

(> 99% confidence level) between the three meteorological parameters and MARdet exists for most (82%) of the data points (shaded area in figure 4.7b). From the running multiple regression analysis (figure 4.7b), it emerges that on average 74% of the variability in the MARdet signal can be explained by variations in the three meteoro- Chapter 4 53 Comparison with climate data logical parameters (table 4.2). The subsequent stepwise removal of parameters decreases the mean correlation coefficient to different degrees, depending on which of the three parameters is removed (table 4.2). The largest reduction of the mean correlation coefficient is observed when mean summer temperature is eliminated. From this it can be estimated that, on Table 4.2: Results of the stepwise, running multiple regression average, at least 50% of the cor- analysis. T=mean summer temperature, S=number of days with snow, P=cumulative summer precipitation, M=detrended, normal- relation can be ascribed to ized mass accumulation rate and r=correlation coefficient. variations in mean summer r2 TSPM TSM TPM SPM TM SM PM temperature, whereas the rest Mean 0,74 0,63 0,61 0,45 0,45 0,30 0,19 can be divided equally be- Minimum 0,10 0,03 0,02 0,00 0,00 0,00 0,00 tween the two other variables. Maximum 0,99 0,97 0,98 0,96 0,97 0,88 0,80 Consequently, it is apparent that summer precipitation and the number of days with snow have an influence on MARdet variations that can not be neglected. Nevertheless, one has to consider that the number of days with snow is correlated with mean summer temperature (r2 = 0.23), so that part of the contribution of this parameter to the signal must in fact be ascribed to mean summer temperature. An example of rainfall influence rather than increased glacial melting is given by the high MARdet values around 1890 (figures 6 and 7a). A striking feature in figure 4.7b is the time interval between 1922 and 1936 that shows no or only a very weak significant correlation with any of the three climatic variables. During this time interval, the three parameters do not fluctuate significantly, but remain close to their long-term mean, as does MARdet (figure 4.6). This might be a sign of a more maritime climate with decreased seasonality (smaller annual differences in temperature and precipitation), as reported by Thompson (1995) for Europe in general and Pfister (1995) for Switzerland in particular, for the 1920s. While the mean of the running correlation coefficients helped to identify the relative importance of the three parameters, the wide range of correlation coefficients (table 4.2) demonstrates the limited use of the mean correlation coefficients. A quan- tification of the contribution from the different parameters to the MARdet signal, aim- ing at the development of a transfer function, is probably not admissible, since their importance changes dramatically through time.

4.4 Conclusions

The sediments of Lej da Silvaplauna deposited during the last 130 years exhibit increasing fluxes of biogenic silica, total organic carbon, total nitrogen and total sul- Comparison with climate data 54 Chapter 4 fur since 1960. A comparison with tourism data (Sils, 1993/94; Silvaplana, 1993/94) suggests that increased anthropogenic nutrient input (causing elevated autochthonous production) due to increased tourism is the reason for the recent eutrophication trend of the lake. Autochthonous components contribute around 10% to MARtot per year since 1960 but are negligible before this date.

After subtracting out the autochthonous components, MARall fluctuations in Lej da Silvaplauna integrate signals for both century-scale and decadal-scale environmental changes that can be separately traced. Low-frequency (decadal to centennial)

MARall variations reflect changes in glacier size. High MARall values reflect large glacier sizes and generally low summer temperatures. The catchment glacier lag time, as estimated from the MARall signal, is very short (less than 4 years) and in agreement with what has been observed for other small (< 3 km2) valley glaciers

(Gamper and Suter, 1978). High-frequency (yearly) MARdet variations represent changes in the meteorological conditions. In contrast, high MARdet values on a short timescale reflect increased glacier melting and therefore (1) generally high summer temperatures, (2) fewer days with snow and/or (3) increased amounts of rainfall.

Extreme conditions are recorded especially well in the MARdet record. According to a running multiple regression analysis, the three climate variables (mean summer temperature, cumulative summer precipitation and the number of days with snow per year) together explain 74% of the variance in the MARdet signal. The stepwise elimination of parameters suggests that on average at least 50% of the correlation can be ascribed to variations in mean summer temperature, followed by summer precipitation and the number of days with snow. Although the regression analysis of the data set yields the above ranking of parameters, it also becomes obvious that the wide range of occurring correlation coefficients prohibits attempts to quantify the different forcing factors.

Finding a method to trace the varying contribution of rainfall to MARdet of each year is important for developing a more quantitative interpretation of the varve record. Once this can be established, it might be possible to develop transfer functions and link sediment signatures like the 1987 flood-turbidite to distinct weather patterns that are known to be characteristic for a heavy rainfall event (e.g. August 1987; Bundesamt für Wasserwirtschaft, 1991). Our “reading” of the pre-instrumental varve record would then be much more refined. From Lej da Silvaplauna and adjacent Lej da Champfèr, varved sediments are available for the last 3500 years (Leemann and Niessen, 1994b) and from about 13000 BP to 9500 BP (Ohlendorf et al., 1995). Hence, there is an opportunity of interpreting the drastic Late Glacial climatic changes in a process-related way. Chapter 5 55 Weathering tracers

CHAPTER 5

Evaluation of sedimentary tracers for catchment weathering and erosion modes

5.1 Introduction

Precise estimates of continental weathering rates are crucial input parameters for global climate models and thus for the prediction of future climate changes, for they play a major role in the global carbon cycle (Walling, 1987) and therefore affect the atmospheric CO2 budget. In particular the need for quantitative estimates of chemical weathering rates especially in glacial environments has recently been stressed by Föllmi (1995); Sharp et al. (1995) and Velbel (1993). Considering the fact that twice the amount of atmospheric CO2 is consumed if silicate weathering dominates over carbonate weathering (e.g. Berner and Lasaga, 1989; Lins, 1994) the high efficacy of chemical weathering in glacial environments (Brown et al., 1996) may play an important role in global carbon cycling on glacial-interglacial timescales (Sharp et al., 1995) that previously has been underestimated (Berner, 1990). In the attempt to obtain quantitative information about the relative importance of different erosional processes, case studies in small scale erosional systems with relatively well defined boundary conditions are an important source of information. Following this pathway, the common objective of all investigations that are summarized in this chapter, was to identify sedimentary proxy indicators that enable us to separately trace the influence of different climatic forcing factors on the annually laminated sediments of Lej da Silvaplauna. It has been shown in chapter 4 that in particular rainfall, through its potential to increase soil erosion, can exert a discernable influence on annual mass accumulation rates in Lej da Silvaplauna (Ohlendorf et al., 1997). This influence currently can not be distinguished from the temperature controlled input of glacigenic rockflour. Therefore the necessity arose to look for a tracer that allows to quantify the influence of soil erosion - considered here as a synonym for rainfall - on annual MAR. The final goal here must be to determine this influence on MAR ideally for every single varve or at least for a group of 5 varves. Conse- quently the desirable tracer has to be of such a nature that it can realisticly be mea- Weathering tracers 56 Chapter 5 sured with such a high resolution. An overview of the use of lake sediment characteristics as tracers for changes in catchment erosion modes is given by e.g. Dearing (1991); Oldfield and Clark (1990) and Thomas (1988). For studies in recent sediments the 1986 fallout of 137Cs from the Chernobyl acci- dent might offer an opportunity to descriminate material transported by meltwater from sub- or englacial sources from material derived from the fluvial erosion of proglacial deposits (Hasholt, 1992). This tracer only is applicable to sediments younger than 1986 and therefore was not considered in this study. In this study 3 different methods were explored for their capacity to reach this goal. Results of the following investigations are presented: (1) Mineralogical analysis of turbidites. The mineralogical signature of the catchment rocks and soils is detectable in the river sediments and should be best preserved in turbidites. (2) Geochemichal composition of selected sediment samples from contrasting lithologic units. It is expected that during times of small or absent catchment glaciation, chemical weathering of catchment soils is intensified. Therefore, sediments deposited under non-glacial catchment conditions, like the early and middle Holocene in Lej da Silvaplauna, should show a different geochemical signature than purely glacigenic sediments. (3) Analysis of magnetic properties. More stable catchment soils during glacier free periods, should give rise to the formation of new magnetic minerals in soils or lake sediments. The potential of these three parameters as rainfall-tracers is evaluated and an assessment about their applicability in high resolution studies is given. For points (1) and (2) prior to the analysis of mineralogical and geochemical changes back in time, a mapping of their todays spatial distribution in the lake sediments and their input by rivers was carried out (5.2.1 and 5.2.3).

5.2 Results

5.2.1 Mineral composition of river- and most recent lake sediments

In the catchment of the Upper Engadine lakes, numerous areas with different lithologies can be distinguished as described in section 2.2 (figure 2.3). In particular the band of ophiolites that is offset by the Engadine line shows a mineral assemblage distinctly different from all other rocks outcroping in the catchment. Consequently the bedload of rivers draining the respective sub-catchments should mirror their lithologies. If this is the case the question arises whether this signal is also transfered to the lake sediments and if temporal variations in the mineralogy of lake sediments can yield information about changes in catchment dynamics through time. Chapter 5 57 Weathering tracers

The mineral composition of river- and most recent lake sediments is presented in figure 5.1. From the 21 rivers draining into Lej da Segl and Silvaplauna sediment samples were taken at their deltas. Only sediment with a maximum grainsize smaller 2.0 km N 1.0 11 12 0 13 21 14 20 15 19 16 17 18 1 . 10 2 9 8 3 4 5 Serpentine Talc Amphibol 7 Chlorite Mica 6 Other Calcite Dolomite Orthoclase Plagioclase ø = Quartz content

Figure 5.1: Mineral composition of river and lake sediments in the catchment of lake Sils and Silvaplana. The diameter of the pie charts represent the quartz content of the sediments. Note the smaller quartz intensities of the lake sediments with respect to the river sediments and the increased serpentine intensities in rivers 8, 9 and 13-15 that are draining catchments with ophiolitic rocks (cf. figures 2.3 and 7.14). Weathering tracers 58 Chapter 5 than 2 mm was sampled to account for the fact that only clay and silt or sometimes sand, in turbidites, are deposited in the profundal zone of the lakes. In addition, from every core taken in the lakes the uppermost one centimeter of sediment, comprising the last 3 to 10 years, was analyzed for its mineral composition, in order to get a picture of the spatial distribution of sediment in the lakes. Because quartz is the dominating mineral (>50%) in all samples, quartz contents do not appear as pie pieces in figure 5.1, but are represented by the diameter of the respective pies. The geologic units referred to in this chapter were described in section 2.2 and presented in figure 2.3.

River sediments The riverdelta sediments clearly reflect the lithology of the catchment they are draining. Rivers draining the northwestern catchment that is composed of granitic rocks from the Julier nappe can be identified by their higher orthoclase and plagioclase contents. This is not so obvious for 2 small rivers at the southern end of Lej da Silvaplauna (19 and 20 in figure 5.1). They are characterized by increased amphibol and dolomite contents instead, because these two rivers drain a sub catchment where diorites and gabbros, which are partly covered by moraines of the engadine valley glacier, are exposed. The two glacial rivers (Aua da Fedoz and Fedacla) that drain the largest parts of the southeastern catchment (Val Fex and Val Fedoz) and also supply the largest amount of suspension show a relatively similar mineral assemblage. The dominating mineral is plagioclase followed by mica and chlorite. Only minor amounts of orthoclase, amphibols and pyroxene occur. The only difference between them is the higher chlorite content of the Fedacla, probably due to the ‘Fexer Schuppenzone’ outcropping in Val Fex. Also the small rivers draining into the southern basin of Lej da Segl show more or less this mineral assemblage representative for the ‘Margna rocks’. it The river that comes from the small cirque glacier (Vadrec da Murtairac) at the northern flank of Piz de la Margna has a slightly higher amphibol content, due to the amphibolites that are exposed there. The quartz contents of the river sediments are highest for rivers draining the southwestern catchment and also for river 15. A striking feature is the high serpentine content of the Ova dal Mulin, Lavatera, Ova da la Roda and the Ova da la Rabquisa + two smaller rivers on the southeastern shore of Lej da Silvaplauna. These rivers drain areas, where the ophiolites of the Upper Pennine Platta nappe outcrop (cf. figure 2.3). Small amounts of calcite, talc and stilpnomelane were detected in the Ova da Surlej, Ova da l’Alp, river 15 and river 5. Although their concentration is rather small, they are very indicative for their source region because their occurence is restricted to a Chapter 5 59 Weathering tracers small area. River 5 is the only one that contains stilpnomelane which occurs in the metasediments enveloping the Maloja-Gneiss (Finger et al., 1982). Talc occurs in the delta sediments of rivers 5, Ova da l’Alp and 15. Talc bearing rocks have been described in the catchment of these rivers where Platta- and Forno-Lizun Ophiolites interfinger with mesozoic sediments (Finger et al., 1982). The Ova da Surlej is the only river that contains noteworthy amounts of calcite because large parts of its catchment are composed of Austroalpine sediments that contain bands of marble (Streifenmarmor, Staub, 1946) were the calcite originates from.

Most recent lake sediments For reasons of readability in figure 5.1 the mineral composition of selected samples from the most recent lake sediments is plotted. The complete sample table can be found in appendix 5.1. In general quartz contents of the lake sediments are lower than those of the river sediments. In the lakes quartz contents increase from the soutwestern basin of Lej da Segl downstream to Lej da Silvaplauna in the northeast. The mineral composition of ‘shelf and slope’ sediments in the lakes clearly is controlled by the river inflow clos- est to the respective sampling site. This is best demonstrated by the very high orthoclase content in core SS 92/3 which is close to the Ova da la Roda and the Ova dal Crot. In contrast core SS 92/6, which is from the profundal zone and also close to these rivers, does not show this feature. Instead, average orthoclase and higher amphibol contents are observed. Since amphibols are only observed in rivers draining the southeastern catchment the profundal sediments seem to integrate the influence of all major rivers feeding into the lakes. The same phenomenon is observed in Lej da Silvaplauna where core SSP 92/7 which is close to the Ova da l’Alp contains serpentine, but core PSP 90/2 does not. In general the content of phyllosilicates is slightly higher in the lake sediments then in the rivers.

5.2.2 Mineral composition of turbidites in Lej da Silvaplauna Holocene sediments

The mineral composition of 54 turbidite layers in the Holocene sediments of Lej da Silvaplauna is shown in figure 5.2. The chronology of Lej da Silvaplauna sediments is well established (13 AMS-dates on macro remains) and was taken from Leemann (1993). Turbidite mineralogy of the lower part of the core, comprising the early and middle Holocene can be subdivided in two groups. The early Holocene turbidites show high plagioclase and orthoclase contents and a relatively low chlorite content. Most of the middle Holocene turbidites have higher chlorite and mica, Weathering tracers 60 Chapter 5 depth PSP Age 0 102030405060708090100 90/1 unit BP (cm) 0 0 1 Legend

50 Talc 100 100 385 150 Serpentine 200 200

250 Amphibole

300 925 300 Chlorite 1015 350

2 400 400 Mica

1725 450 Plagioclase 500 500

550 Orthoclase 600 600 2775 2875 650 Dolomite

700 700 2965 Calcite 750 3835

800 800 Other 4620 850 5845 900 7830 900 Pyroxene 3 8755 9100

Eastern catchment with mesozoic sediments Ova da Surlej (11) Southeastern, ophiolithic catchment River 15 Southeastern, glaciated catchment Fedacla (16) Northwestern, granitic catchment Ova dal Valhun (21)

Figure 5.2: Mineral composition of turbidites in the Holocene sediment sequence of lake Silvaplana (core PSP 90/1 from Leemann, 1993). Relative mineral percentages not considering quartz intensities are poltted. At the bottom mineral compositions of 4 rivers that are characteristic for the lithologies of the different catchements are shown. but lower plagioclase and orthoclase contents. Two turbidites, one at the beginning and the other one at the end of the middle Holocene, contain small amounts of calcite. The turbidites bracketed by these two calcite-bearing turbidites show traces of serpentine, talc or stilpnomelane. During the late Holocene turbidite composition changes several times between two different types. One type with high plagioclase, Chapter 5 61 Weathering tracers orthoclase and dolomite content and another type with high mica contents. Among the latter type a few samples contain small amounts of talc, serpentine or stilpnomelane.

5.2.3 Geochemistry of the most recent and Holocene lake sediments Most recent lake sediments The main and trace element composition was determined in 24 selected sediment samples taken from the recent (uppermost 1.0 cm) and subrecent sediments of Lej da Segl, Silvaplauna and Champfèr. Since sediment structure varies spatially, not all of the sediments are annually laminated. This precludes the precise calculation of element-accumulation rates. The results summarized in figure 5.3 therefore are given as element concentrations. As reference values the average element concentrations for shales, granites and basalts after Turekian and Wedepohl (1961) are indicated by arrows. Most of the measured values fall in the range given by these 3 reference data-sets. However, a detailed comparison of all measured samples reveals several striking features. 1. Except for CF 93/2, all samples show Al-concentrations higher than any of the reference data-sets. This is also true for most of the Mn-concentrations of the recent sediments. 2. Two sediment ‘profiles’ in Lej da Silvaplauna (SSP 93/1, 0-1 cm and 38-39 cm; SSP 90/3 0-1 cm and 80-81 cm) show increasing concentrations of redoxsensitive elements (Fe, Mn) towards the sediment water interface. 3. Lej da Champfèr profundal sediments (CF 93/2) in general show very low concentrations for most main and trace elements except for Zn were the highest concentration of all samples was measured in Lej da Champfèr. 4. Sample SSP 93/1 shows by far the highest Cu values of all samples. 5. Sample SS 92/6 also has increased Cu values and in addition shows the highest Na, K and Al but lowest Mn concentrations.

Holocene sediments Element concentrations were also determined in 5 mix-samples from each lithological unit and 3 representative spot-samples from unit 1, 2 and 3 of Lej da Silvaplauna Holocene sediments (figure 5.4). Although most of the concentrations determined in the mix- as well as in the spot-samples are in the range of the 3 reference data-sets there are some exceptions to this: Al-concentrations of samples 2amix -

4mix as well as 1spot - 2spot are higher than any of the reference data sets. The same is true for Mn-concentrations of samples 1mix - 3mix as well as 2spot and the Zn-concentration of sample 2bmix as well as 2spot. In detail the temporal trends in element concentrations can be described as follows. Na and Mg concentrations progressively de- Weathering tracers 62 Chapter 5 ] kg / B B mg [ S S Cu Cu [mg/kg] G G 0 4080120 0 4080120 ] kg / B B S S mg [ Zn Zn [mg/kg] (39) (39) G G 40 60 100 140 40 60 100 140 S S (465) (465) ] B B kg / mg [ Sr Sr [mg/kg] G G 50 100 200 300 50 100 200 300 ] 6000 kg / mg [ Mn Mn [mg/kg] BS BS 0 3000 0 3000 6000 G G (8.65) (8.65) B B Fe [%] Fe [%] S S (1.42) (1.42) 4.0 6.0 8.0 4.0 6.0 8.0 G G [%] Na Na [%] G G B B S S 0.0 4.0 8.0 12.0 0.0 4.0 8.0 12.0 G G K [%] K [%] S S (0.83) (0.83) B B 2.0 4.0 6.0 2.0 4.0 6.0 [%] S S 10.0 14.0 Al [%] Al B B G G 6.0 6.0 10.0 14.0 (4.60) (4.60) B B [%] Mg Mg [%] S S (0.16) (0.16) G G 1.0 3.0 5.0 1.0 3.0 5.0 5.0 (7.60) (7.60) B B S S Ca [%] Ca [%] G G 0.0 2.5 0.0 2.5 5.0 SS 92/2 SS 92/6 SS 91/2 SS 91/1 SS 93/1 SS 93/4 SS 93/2 SS 93/3 CF 93/2 SSP 93/5 PSP 90/2 PSP 90/3 SSP 93/4 SSP 93/1 SSP 93/2 SS 92/2 7-8 SS 91/2 7-8 SS 93/1 4-5 SS 91/1 4-5 SS 92/2 11-12 SS 91/2 12-13 SS 93/1 10-11 SSP 93/1 38-39 PSP 90/3 80-81 . 2.0 km N 1.0 0

Figure 5.3: Element concentrations in selected samples from the most recent sediments of lake Sils, Silvaplana and Champfèr. For some samples short element profiles are shown. For comparison reference values from Turekian and Wedepohl (1961) are indicated by arrows at the bottom of the diagrams. B= average Basalt, S=average shale, G=average Granite. Chapter 5 63 Weathering tracers

Lake Silvaplana Na [%] K [%] Ca [%] Mg [%] Al [%] MAR [mg/cm2∗yr] Depth PSP Age Sam- 1.00 1.20 1.40 1.60 2.00 2.50 3.00 3.50 4.00 1.00 1.50 2.00 2.50 2.00 3.00 4.00 7.00 7.50 8.00 8.50 9.00 9.50 0 100 200 300 (cm) 90/1 BP ple Unit 0 1 1 10.54 1

100 385

200 2a

300

925 400 2 1015

9.60 500 1725

2b 600

700 2775 2875 2965 800

3835

900

3 4620 3 5845 1000 7830 8755

1100 9415 4 4

S(0.96) G(2.58) B(0.83) G(4.20) G(0.51) B(7.60) G(0.16) B(4.60) S S GBS B(1.80) S(1.50)

Lake Silvaplana Sr [mg/kg] Cu [mg/kg] Zn [mg/kg] Mn [mg/kg] Fe [%] Fe/Mn Depth PSP Age Sam- 100 110 120 130 140 0 5 10 15 20 50 100 150 200 1000 1250 1500 1750 2000 4.00 5.00 6.00 24 26 28 30 32 34 36 38 (cm) 90/1 BP ple Unit 0 1 1 146 1

100 385

200 2a

300

925 400 2 1015

6.43 500 1725

2b 600

700 2775 2875 2965 800

3835

900 157 3024

3 4620 3 5845 1000 7830 8755

1100 9415 4 4

S(300) S(45) G(39) S(850) G(1.42) B(8.65) G G S B B S B(465) B(87) G(390)

Figure 5.4: Element concentrations in the Holocene sediments of lake Silvaplana in mix (black areas) and spot (small white bars) samples. Note high Zn and Ca concentrations in sample 2b and 3 respectively and low K, Al caoncentrations and Fe/Mn-ratios in sample 3. For comparison reference values from Turekian and Wedepohl (1961) are indicated by arrows at the bottom of the diagrams. B=average Basalt, S=average shale, G=average Granite. Weathering tracers 64 Chapter 5

crease towards younger sediments from sample 4mix to 1mix. The opposite trend is observed for the spot-samples. Concentrations of K, Al and Fe are low in sample 1mix and 3mix and highest in samples 2a+bmix. In contrast in the spot-samples, K, Al and Fe concentrations are high in the youngest sediments (sample 1spot) but still lowest in sample 3spot. Cu concentrations show only very little variation in both the mix- and the spot-samples. The same holds for Zn except that very high Zn concentrations occur in sample 2bmix and 2spot. In mix- as well as spot-samples Ca and Mn concentrations show maximum values in sample 1mix, 3mix, 1spot and 3spot. The same is true for Sr concentrations in the spot-samples, whereas the mix-samples show low values in sample 1mix. For these samples average element accumulation rates were calculated based on the chronology presented in Leemann (1993) (appendix 5.2). Bulk sediment accumulation on average is lower by a factor of 6.8 in the early and mid Holocene sediments (sample 3) compared to the time periods before and after (appendix 5.2). The fact that element accumulation rates are lowest in this time intervall too, indicates that they are mainly controlled by variations in the amount of allochthonuous minerogenic input rather than by changes in the limnological conditions, like bottom water redox status, or input from atmospheric and anthropogenic sources. For this reason element concentrations should -150 -125 -100 -75 -50 -25 0 25 50 75 mirror changes in the compo-

Na glacial/recent basin plain sition of allochthonous matter. glacial/recent base of slope K glacial/non-glacial However, if an element en- Ca richment factor deviates from Mg the value anticipated from the Al Depletion Enrichment difference in bulk sediment ac- Fe cumulation (e.g. 6.8 for sample

Mn 2bmix vs. 3mix), then the ob-

Sr served element enrichment or

Zn depletion is most likely caused

Cu by atmospheric input or changes in the limnological Figure 5.5: Element enrichment or depletion factors of A basin plain, base of slope and non-glacial with respect to glacial sediments are conditions. In figure 5.5 devia- shown. A strong depletion of Zn and Cu is observed in all 3 samples. tions from the anticipated fac- Note that Mn is enriched in the recent basin plain and non-glacial sediments, but not in the base of slope sediments. tor are plotted for spot-samples

3spot and 1spot as well as sample

PSP 90/3-1. Reference value is always spot-sample 2spot, which represents sediments deposited under glacial conditions. Sample 3spot represents sediments deposited under non-glacial conditions. Sample 1spot (PSP 90/2) represents recent base of slope Chapter 5 65 Weathering tracers sediments, whereas sample PSP 90/3-1 represents recent sediments of the basin plain. The non-glacial sediments show an enrichment in Ca, Mn and Sr and a depletion in K, Al, Fe, Zn and Cu. Mn enrichment and Zn depletion are particularly strong. The recent basin sediments are also enriched in Ca, Mn and Sr and also slightly in Na, Al and Mg. They are depleted in Cu and Zn. The same pattern is observed for the recent base of slope sediments with the marked difference, that they are not enriched in Mn and slightly depleted in Fe.

5.2.4 Magnetic properties of Holocene sediments in Lej da Silvaplauna

157 samples of Lej da Silvaplauna Holocene sediments (core PSP 90/2) were subject to frequency-dependent susceptibility (cfd) measurements (3.4.1). The results are plotted in figure 5.6a as F-factor changes through time. F-factor variations should yield information about the contribution of fine magnetic mineral grains to the susceptibility signal. The chronology was adapted from Leemann and Niessen (1994b). Susceptibility of Lej da Silvaplauna sediments in general is relatively low. For the HF/LF-measurements values between 0.9 and 26.2 * 10-6 were determined. When F- factors are calculated, 28 of 157 samples yield negative values. Susceptibility at high frequencies should always be lower or at most equal to that measured at low frequencies, depending on the amount of magnetic grains with a grainsize at the superparamagnetic/stable single domain border that are present in the sample. Con- sequently negative values should not occur. They therefore might be an effect of the low overall values that are close to the precision of the measurement and are disregarded for the interpretation. For the same reason also the other values have to be interpreted with caution. Values subsequently referred to as ‘high’ are greater than one standard deviation and are considered to represent a ‘true’ increase in F-factor. In figure 5.6b only these ‘high’ values are potted. Based on this partition, the following general trends can be deduced from the dataset. Sections with prevailing positive F-factors were detected at 0-117, 159-455, 655- 1140 cm. High F-factors occur in the early Holocene (at 6700 and at 7400 BP), between 3900 and 5200 BP, between 2600 and 3400 BP, at the transition from the middle to the late Holocene, in the late Holocene (from 1500 to 700 BP) and in the subrecent sediments (250 to 0 BP). The highest F-factor was measured at 56 cm sediment-depth (170 BP = AD 1820). Weathering tracers 66 Chapter 5

F-factor [%] F-factor > 1Θ [%] Depth PSP Age -8 -4 04812 0246810 12

(cm) 90/1 Unit BP 0 1

28 28 100 385

200

300 925 1015

400 2

500 1725

600

700 2775 2875 2965 800 3835

900

3 4620 5845 1000 7830 8755

1100 9415 4 a) b)

Figure 5.6: F-factor (frequency dependent susceptibility) changes in the Holocene sediments of lake Silvaplana. a) F-factor rawdata b)F-factors that are greater than 1 standard deviation. Chapter 5 67 Weathering tracers

5.3 Discussion

5.3.1 Recent controls on sediment distribution

It has been demonstrated (5.2.1), that the composition of river-delta sediments is a reflection of the sub-catchment lithologies. River sediments thus are regarded as fingerprints of their particular catchment. But the mineralogical analysis of the most recent lake sediments (5.2.1) revealed that mixing of the different lithologic signatures occurs in the lakes with increasing distance from the deltas. As a consequence the mineralogic ‘fingerprints’ of the different catchments are blurred in the basin plain sediments. However, one feature that came out of the analysis of the most recent lake sediments (figure 5.1) is the fact that lowest quartz contents are observed in the southwestern basin of Lej da Segl, which presently does not receive glacial meltwaters, whereas higher quartz contents occur in front of the mouth of the two glacial rivers. This expresses the importance of glacial erosion as a sediment source. Only the grinding process at the base of a temperate glacier is capable of producing large quantities of rockflour with a high quartz content also in the fine fraction and a very small mean grain size (Small, 1987a). Transported in the glacial stream as glacial milk and finally injected into the stratified watercolumn of a downstream lake as under- and interflow (Sturm, 1979; Sturm and Matter, 1978) this rockflour leaves behind several characteristic fingerprints in the sediments of the receptive lakes. One of these fingerprints is the occurence of fine grained quartz, others are discussed in chapter 7. It seems that all of the fine grained quartz is trapped in the lake basins that directly receive the glacial meltwaters i.e. Lej da Segl northeastern basin and Lej da Silvaplauna. This is suspected because Lej da Segl southwestern basin and in particular Lej da Champfèr sediments show the lowest quartz contents of all samples. Even the quartz-rich rivers that drain into the southwestern basin of Lej da Segl do not leave a signature, because they are not glacial rivers and therefore have low flux rates and do not contain finegrained quartz. In Lej da Champfèr X-ray amorphous autochthonous components like organic matter and biogenic silica make a major part of the sediment. This might be the reason why intensities in the diffractogramm, which represent the detrital mineral input, are very low.

5.3.2 Turbidite mineralogy: Tracer for catchment drainage conditions ?

To avoid the problem of a blurred mineralogical signature the mineralogy of turbidites occuring in the sediment section was determined. Their mineral composition should reflect that of their source region most directly since it is not altered by mixing processes within the lake. In particular the following mineral assemblages should Weathering tracers 68 Chapter 5 be distinguishable (cf. figure 5.2). 1. Turbidites from the glaciated catchment of the Fedacla should contain quartz, plagioclase, mica and chlorite, but no or only very little orthoclase, amphibols and pyroxene. 2. If, in addition to the assemblage described under 1.), serpentine or amphibol or talc are also present the origin of the turbidite most probably is the delta of the Ova da la Rabquisa or the two smaller adjacent rivers (13 and 15). 3. Turbidites from the non-glaciated northwestern catchment (in particular Ova dal Valhun) should predominatly consist of quartz and plagioclase and also contain orthoclase. 4. If calcite occurs in a turbidite the most likely source is the delta of the Ova da Surlej. It is expected that a change in the erosional regime from a non-glacial i.e. probably rainfall dominated, to a glacier dominated system as it took place in the catchment of Lej da Silvaplauna at the transition from the middle to the late Holocene, would also result in an activation of different sediment sources. If this is true the mineral assemblage of turbidites that were released during non glacial times should differ from those released in glacial times. A change in the source region of turbidites in turn could be an indicator of changed catchment dynamics eventually due to changed atmospheric circulation or increased melting of permafrost at certain places. From this it seems feasible to infer the source regions of turbidites observed in Lej da Silvaplauna profundal sediments (core PSP 90/1) from their mineralogical ‘fingerprint’. However, for the bulk mineral composition of turbidites presented in figure 5.2, this assumption is misleading because it ignores the sorting process that is an inherent feature of turbidite deposition. Therefore, (1) the effect of sorting has to be considered in order to (2) understand how the mineralogical signal of the source region is preserved in turbidites, (3) which major signals can be distinguished and (4) what they means in terms of catchment dynamics.

The effect of sorting When a turbidity current is released, under waning flow regime of the depositing turbidity current in general fining occurs in a downcurrent as well as in an upward direction (Bouma, 1962; Rupke, 1978). The fining upward that is observed in a turbidite layer also controls the mineral distribution. This becomes obvious when in the turbidites from PSP 90/1 the basal sandy silt layers and the clay-tops are analyzed separately. Isometrical and stalked minerals like feldspars, dolomite and amphibol are enriched at the base whereas the platy minerals chlorite and mica occur at the top. The same separation is observed from the proximal to the distal parts of a turbidite. Consequently the bulk mineral assemblage of a turbidite in core PSP 90/1 in the first place yields information about its distance from the source region, rather than about the mineral composition of the source itshelf. Considering this, for core PSP 90/1, which was taken at the northeastern end of Lej da Silvaplauna, a high Chapter 5 69 Weathering tracers [%] 4 Other Calcite Stilpnomelane Serpentine Talc 02 68 0.2 Amphibole / Plag. Amphibole 0.0 0.1 0.3 0.4 1.0 Dolomite / Plag. Dolomite 0.0 0.5 1.5 2.0 Orthoclase / Plag. Orthoclase 0.0 0.2 0.4 0.6 0.8 Chlorite / Mica Chlorite / 0.5 1.5 2.5 3.5 [%] a) b) c) d) e) f) 30 50 70 90 Chlorite+Mica Chlorite+Mica 0

1000 2000 3000 4000 5000 6000 7000 8000 9000

C-years BP C-years 10000 14 925 385 BP

Age

2775 2875 2965 3835 4620 5845 7830 8755 9415 1015 1725 Variations in mineral ratios and occurence of trace minerals in the Holocene sediments of lake Silvaplana between 0 and 10000 BP. a) Chlorite+mica, b) Chorite/ a) Chlorite+mica, between 0 and 10000 of trace minerals in the Holocene sediments of lake Silvaplana BP. in mineral ratios and occurence Variations Unit 1 3 4 2 PSP 90/1 0 700 800 900 500 600 200 300 400 100 1000 1100 Lake Silvaplana (cm) Depth Figure 5.7: Figure of stilpnomelane, calcite, serpentine and talc. contents d) Dolomite/Plagioclase-ratio, e) Amphibole/Plagioclase ratio, f) relative Mica-ratio, c) orthoclase/plagiclase-ratio, Black dots indicate sample positions. Weathering tracers 70 Chapter 5 percentage of phyllosilicates (chlorite + mica) would be indicative of turbidites from the delta of the glacial river (Fedacla), because they have the longest travel distance. This is only true however, when other effects that could control the transport compe- tence of a turbidity current like the lake bottom topography or pre-sorting of the sediment are disregarded. Chlorite plus mica percentages are plotted in figure 5.7a.

How is the signature of the source region preserved ? Judging from the calculated mean annual discharge (LIMNEX, 1994), the most probable other turbidite sources, after the Fedacla and Sela with 1.498+1.989 m3/s discharge, are the Ova dal Valhun and the Ova da Surlej with 0.705 and 0.327 m3/s, respectively. These two rivers are equally close to the coring site. A further determination of turbidite sources therefore required a separate analysis of the composition of platy minerals on the one hand and isometrical/stalked minerals on the other hand in order to exclude the sort- 6 0.4 ing effect. The signature of the 0.35 5 Chl/Mica source region consequently is Dol/Plag 0.3 Amph/Plag 4 preserved in the mineral ratios of Kf/Plag 0.25 a turbidite. Thus mineral assem- 3 0.2 blages are expressed as chlorite/ 0.15 2 mica- and orthoclase, dolomite 0.1 1 0.05 or amphibol/plagioclase-ratios

0 0 in figure 5.7b-e. For comparison 11 12 13 14 15 16 17 18 19 20 21 the same ratios for the 11 rivers Sela Fedacla entering Lej da Silvaplauna are Ova da SurlejOva da l'Alp Ova dal Valhun Ova da la Rabqiusa given in figure 5.8. In addition,

Figure 5.8: Mineral ratios (chlorite/mica, orthoclase/plagioclase, indicator minerals like talc, ser- amphibole/plagioclase and dolomite/plagioclase) in the sediments pentine, calcite and stilpno- of rivers draining into lake Silvaplana (11 to 21 in figure 2.6). melane (figure 5.7f) can help to identify the turbidite source regions, even though they usually occur in only small amounts.

The two main mineralogical signatures The two major mineralogical signatures that allow to infer changes in catchment drainage conditions are the ‘glacial’ and the ‘non-glacial’ signature. The ‘glacial signature’ means turbidites that originate from the Fedacla delta. The Fedacla is the only river that drains a glaciated catchment. Turbidite generation at the Fedacla delta therefore should represent increased sediment production in the southeastern catchment and is taken as evidence for glaciation. Fedacla turbidites, i.e. the ‘glacial sig- Chapter 5 71 Weathering tracers nature’, can be recognized by high phyllosilicate percentages and low chlorite/mica ratios. The ‘non-glacial signature’ represents turbidites from the northwestern and the northeastern catchment (Ova dal Valhun and Ova da Surlej) were no glaciers are present. This signature is characterized by low phyllosilicate percentages and high chlorite/mica ratios. It prevails when the otherwhise dominating influence from the Fedacla is reduced or not present. This is the case when glaciers in the southeastern catchment are small or not present like in the early and middle Holocene.

Climatic interpretation of the turbidite record When the mechanisms described above are taken into account the following interpretations can be made from the turbidite record. During the time intervals 9000 to at least 6200, 4700-4500, 3800-3050, 2900-2650, 2000-1600 BP turbidites with low phyllosilicate percentages and high chlorite/mica ratios occur frequently. In contrast, during the time intervals 2650-2000 and 1600-900 the majority of turbidites shows high phyllosilicate percentages and low chlorite/mica ratios. This change in turbidite characteristics indicates a change in the source region from the northwestern to the southeastern catchment. According to the above explanation this represents the change from a ‘non-glacial’ to a ‘glacial’ mode. Also the first measured turbidite of the profile at 9100 BP can be interpreted as one from the southeastern catchment, based on its phyllosilicate percentage, although the chlorite/mica ratio is not especially low. Apparently, although Leemann and Niessen (1994b) identify advancing glaciers in Val Fex already at 3300 BP from a decrease in grainsize and the onset of varve formation in Lej da Silvaplauna sediments, turbidite generation shifts to the Fedacla delta not before 3000 BP when the first ‘glacial’ turbidite is observed. A clear sustained shift to the ‘glacial’ mode occurs as late as 2650 BP. At this time the beginning of a glacier advance that lasts until 1850 BP, is indicated by increasing accumulation rates (Leemann, 1993). Phyllosilicate contents and chlorite/mica ratios indicate delta activity in the southeastern catchment also from 4400 to 4000 BP. But since the respective turbidites contain talc or serpentine, the Ova da la Rabquisa (or rivers 13 or 15) are the most probable source regions. The same is true for turbidites around 4800, 2450, 1350 and 1150 BP. Among all turbidites with a non-glacial signature only two, at 3800 and 6200 BP, contain calcite and therefore most likely originate from the Ova da Surlej. All other non-glacial turbidites originate from the northwestern catchment, presumably the Ova dal Valhun, as indicated by their high orthoclase/plagioclase ratios. Exception- ally high dolomite/plagioclase ratios of two turbidites at 2790 and 2220 BP point to river 19 or 20 as source region. Weathering tracers 72 Chapter 5

From the Boreal until the middle Atlantic with no or only minor catchment glaciation (as identified from the grainsize record), delta activity apparently was restricted to the northwestern catchment mostly to the Ova dal Valhun. This most likely represents turbidites generated in a rainfall dominated system because the catchment of this river is the biggest one after the - at that time inactive - glacial catchment in the southeast (Fedacla). Today it also shows the second highest discharge values (LIMNEX, 1994). The Ova da Surlej at that time eventually was draining into Lej da Champfèr. In late Atlantic and particulatly in the early Subboreal the occurrence of several turbidites from the southeastern, non-glacial catchment might be interpreted as an indication for decreasing slope stablity in response to disappearing permafrost on the northwest facing slopes of the mountains between Piz Murtel and Piz Chapütschin (above the plateau of Furtschellas). Today the realatively small catchment of these rivers together with their low discharge values document their unimportance as a sediment source. On the other hand significant portions of their catchment are occupied by discontinuous permafrost (Hoelzle and Haeberli, 1995). Spatial modeling of the glacier and permafrost distribution in the Corvatsch- Furtschellas area (Hoelzle, 1994) demonstrated that permafrost could vanish rapidly in large areas under a sustained warming trend. Although a 1°C warming according to IPCC scenario A (Houghton et al., 1990) would result in the degradation of only 4% of the permafrost. 65% of the permafrost would disappear if the temperature increases by 3°C. Thawing of supersaturated permafrost deposits leads to slope instability and the generation of landslides and debris flows (Haeberli et al., 1993). Considering this, delta activity in the southeastern, non-glacial catchment during the late Atlantic to early Subboreal and also during the middle Subatlantic (1350- 1100 BP) might reflect increased geomorphologic activity as a consequence of permafrost degradation. In contrast, during Neoglacial times the delta of the Fedacla became active again in response to glacial activity, but with a lag time of several hundered years. Hence from the middle Subboreal to the beginning of the Subatlantic the northwestern catchment again was the dominating sediment source before turbidite generation switched to the southeastern deltas. The lag time probably occurs due to the fact that the advancing glaciers in Val Fex do not release a lot of sediment and that proglacial sediment traps have to be ‘saturated’ before deltaic turbidite generation in response to excessive sediment supply can start again. In summary 3 conclusions may be drawn from the analysis of the Holocene turbidite record in Lej da Silvaplauna. (1) Turbidite generation at the delta of the glacial river Fedacla responds to the transition from non-glacial to Neoglacial catchment conditions around that occured around 3300 BP with a lag time of several hundered years. (2) Turbidites from the southeastern catchment coincide with glacier advances starting around 2650 and 1600 BP. (3) Turbidites from the non-glaciated southeastern Chapter 5 73 Weathering tracers catchment between 4800 BP (max. 6200 BP) and 3800 BP might indicate delta activity due to the degration of permafrost.

5.3.3 Chemical properties: Tracers for the intensity of chemical weathering ?

The geochemistry of the different metamorphic sedimentary and ignous rocks outcroping in the catchment of the lakes covers a wide compositional range. Conse- quently variations in the element composition of Lej da Segl, Silvaplauna and Champfèr clastic sediments are quite large. The comparison with the average composition of different reference rocks given in figure 5.4 illustrates this. Clastic particles from glacial abrasion and soil erosion are by far the most important constitu- ents of the sediments (cf. chapters 4 and 7). Hence element concentrations should approximately fall in the range given by the reference data-sets, if physical weathering is the only factor controlling them. As mentioned above (5.2.3) for a number of samples this is not the case. Therefore changes in chemical and biological conditions of catchment soils and changes in limnological conditions have to be considered as additional factors affecting the spatial and temporal distribution of elements in lake sediments. But not all of the elements are influenced by these different factors in a similar way. Based on their individual behavior in the different environments (soil, lake sediments), groups of elements can be identified that are predominantly affected by only one of those factors. A general assumption is that Na, K and Al are associated with the clastic fraction, Ca, Mg and Sr with the carbonatic fraction, Fe and Mn with the redox sensitive fraction and Zn and Cu with the organic fraction of the sediments. In spite of this complexity chemical signatures of lake sediments can yield information about the erosional history of a basin as has been reviewed by e.g. Dearing (1991) and Thomas (1988). In this section (1) the spatial changes of element concentrations in the most recent lake sediments are discussed. (2) possible signals of changes in the weathering mode are presented. (3) the temporal variations in Lej da Silvaplauna Holocene sediments are interpreted. (4) Complications in the interpretation of the geochemical signals are discussed.

The recent lake sediments Although only a few analysis of the most recent lake sediments were done the different geochemical characters of the lakes are visibile. Lej da Segl and Silvaplauna can mirror changes in the allochthonuous clastic input, whereas in Lej da Champfèr the allochthonuous signal is more suppressed. Weathering tracers 74 Chapter 5

In Lej da Champfèr especially low concentrations of elements associated with the clastic and carbonatic fraction (Na, K, Al, Mg, Ca and Sr) are observed. This most likely is due to the diluting effect of organic matter. The organic carbon content of recent Lej da Champfèr sediments is very high (4-5%). In contrast the concentrations of redox sensitive elements (Mn and Fe) are almost as high as in the other lakes, because they are controlled by lake internal processes. The high Zn concentration in Lej da Champfèr sediments can have several reasons. One might be the enhanced input of waste water from the nearby sewage plant in Surlej that drains into Lej da Champfèr. Municipal waste water according to Förstner (1992) has high Zn-concentrations. Increased immission from the atmosphere is another possible Zn source. For instance Hilgers et al. (1993) and Wedepohl, (1978) showed that soot particles from the combustion of coal and oil have high Zn concentrations too. A third reason might be an enhanced precipitation of Zn from the water. Zn, like most of the other heavy metals, has a strong tendency to form organo-metallic complexes (Dhillon et al., 1975; Förstner and Wittmann, 1979) with humic and fulvic acids. With decreasing Eh-values precipitation of stable Zn sulfide, carbonate or hydroxide is also likely to occur (Förstner and Wittmann, 1979). For Lej da Segl and Silvaplauna erosion controlled sediment sorting is seen by the generally high Al concentrations in all samples. These probably mirror the enrichment of phyllosilicates in the lake sediments with respect to the catchment rocks and/or the erosion of Al enriched surface soils (see discussion of element profiles for further details). In Lej da Segl and Silvaplauna a Zn enrichment towards younger sediments is observed in the sediment profiles (figure 5.3). Therefore an anthropogenic rather than a geologic control for Zn concentrations in the most recent sediments is suggested. An analysis of the spatial distribution of Mn in the surface sediments of Lej da Silvaplauna reveals a weak correlation with the water depth (r2=0.46). However, highest concentrations are not observed at the deepest location, but at core SSP 93/4 which is situatated close to the southeastern shore at the base of a very steep slope. The observed Mn distribution pattern (figure 5.9) might be explained by the circulation characteristics of the lake. Temperature and oxygen profiles measured in May 1994 indicate the developement of an inclined thermocline (figure 5.9). In the middle of the lake the thermocline is encounterd at a waterdepth of 4 m whereas further to the northeast measurements show the thermocline at 19 m waterdepth. This inclina- tion is believed to be caused by the influence of the Malojawind that reaches its maximum strength at Lej da Silvaplauna and pushes its watermasses towards the Chapter 5 75 Weathering tracers

northeast during the days in summer. This phenomenon might also transfer to the redoxcline that developes in the hypolimnic waters. That would result in a better oxygenation of the hypolimnic waters in the northeast and reduced supply in the southwest. Anoxic conditions in the hypolimnion of Lej da Silvaplauna or Segl have never been reported so far (Bosli-Pavoni, 1971; LIMNEX, 1994). However, sedimen-

0 1.0 2.0 km

Mn [mg/kg]

6000 SSP 93/4

PSP 90/3 3000 SSP 93/1 SSP 93/5 Slightly cooler Temp. PSP 90/2 Temp. SSP 93/2 [ C] water of [ C] ° CH 93/2 depth [m] 4° 56 7 56 7 river 11 0 0 10 20 Thermocline 30 measured in may '94 40 50 60 70 75 anticipated chemocline Figure 5.9: Manganese concentrations in the most recent sediments of lake Silvaplana with respect to the position of the thermocline determined in May 1994 by high resolution temperature measurements.

tological evidence - black color up to the sediment water interface, high sulfur accumulation rates (Ohlendorf et al., 1997) - suggests seasonal suboxic conditions in the deepest part of the lake since 1970. If the redoxcline rises above the sediment/water interface during summer, anoxia might be most severe and prolonged in the protected southeastern part of Lej da Silvaplauna also due to the assymetry of the the stratification described above. Under such conditions geochemical focussing, like it has been proposed by Schaller (1996) for Lake Baldegg, might also be an important forcing factor in causing the observed Mn distribution. The reason for the very unusual composition of the surface sediments at sample SS 92/6 remains unclear for the moment. Especially the very high Na concentration (11.94%) is hard to explain. Moreover, also an explanation for the remarkably high Cu concentration of the surface sediments in SSP 93/1 remains to be found. It is supposed that, because sample SS 92/6 in general has a very different element composition - Al, K, Na, Cu values are unusually high whereas Fe and Mn values are unusually low - a drastic change in mineral composition is the most likely reason for Weathering tracers 76 Chapter 5 this. In contrast in sample SSP 93/1 only the Cu concentration is high, which rather points to an anthropogenically caused contamination. The outlet of the sewage plant of Segl-Maria and a small industrial area are close to the location of sample SSP 93/ 1 and must be considered as potential sources of pollutants.

Possible weathering signals It is commonly assumed that elements that might react sensitive to changes in the weathering mode are Na, K, Mg, Ca, Fe and Mn. It has early been recognized by Makereth (1966) that Na, K and also Mg concentrations in lake sediments are to a large extent controlled by the counteracting processes of soil erosion versus soil leaching. Low sedimentary Mg, Na and K values indicate more effective leaching of stable soils, i.e. intensive chemical weathering of feldspars and subsequent removal of cat- ions from the system in solution, when there is sufficient available moisture. In contrast enhanced erosion of barren mineral soils leads to high Mg, Na and K concentrations in the lake sediments. Dean (1993), found that in the sediments of Elk Lake in particular the concentration of Na provides a good index to the degree of chemical weathering in the drainage basin under humid conditions. The author interpretes the presence of a relatively high abundance of plagioclase and high Na concentrations in the sediments of Elk Lake (Minnesota) between 4000 and 9000 varve years BP as an evidence for much drier conditions and less decomposition of feldspars in rock-weathering products in the drainage basin during this time. However, Engstrom and Wright Jr. (1984) emphasized the need to isolate the allogenic phase of the sediments as the carrier of information about the catchment weathering history. The sequential extraction procedure they suggest to accomplish this was not applied in this pilot study. Detected variatons in sediment chemistry therefore can not unequivocally be linked to changes in the catchment weathering mode. Hence, the selectivity of chemical weathering as pointed out by Garrels and Mackenzie (1967) is one of the key features in rock weathering. For instance St. Arnaud and Sudom (1981) describe mineral losses from a luvisolic soil that developed from a glacial till in central Saskatchewan of 40% for Ca-feldspars, 22% for Na-feldspars and 17% for K-feldspars. Calcite and dolomite have been completely weathered from the upper soil horizons. In particular for high elevation catchments in the Rocky Mountains the importance of selective dissolution was underlined by Drever and Hurcomb (1986) and Mast and Drever (1990). They demonstrated that chemical weathering is highly selective since only calcite - which merely occurs in trace amounts - was dissolved from the ignous and metamorphic catchment rocks. The authors at- tribute this phenomenon to the former glaciation of the catchments. With glacier Chapter 5 77 Weathering tracers retreat large amounts of talus and rock flour from glacier erosion were exposed and chemical weathering of these fresh rock surfaces became more important. The fact that rock flour produced by glacial erosion is extremely geochemically reactive has recently been emphasized by Brown et al. (1996). As mentioned above this high efficacy of chemical weathering in glacial environments may play an important role in global carbon cycling that previously has been underestimated (Sharp et al., 1995). A transport of dissolved Ca with subsequent re-precipitation in the lake water therefore seems plausible also in high alpine environments. Furthermore, Ca has a strong affinity for organic ligands such as humic and fulvic acids (Engstrom and Wright Jr., 1984). Organic rich sediments thus may contain substantial amounts of Ca that is not associated with detrital or authigenic minerals. Fe and in particular Mn are elements that react very sensitive to changes in redox- conditions (Stumm and Morgan, 1981). Redox changes can occur either in the lake bottom waters and in the sediments or in the catchment soils. The latter can lead to a reductive dissolution of first Mn and then Fe in the soils. Subsequent transport to the lake occurs in solution, as organic complexes or as colloidal particles (Engstrom and Wright Jr., 1984). According to Makereth (1966), due to the differential mobility of these two elements - Mn is easier released from the soil than Fe because of its greater solubility - the Fe/Mn-ratio can be used as a tracer for changes in soil conditions. Following his interpretation low Fe/Mn-ratios point to reducing conditions in well developed soils, whereas high ratios indicate barren mineral soils. Mobilisation is especially pronounced in soils that develop beneath coniferous vegetation (Engstrom and Wright Jr., 1984) like in the catchment of the Upper Engadine lakes. On the other hand Al and probably also Fe from the decomposition of aluminosilicates can be retained in the catchment soils as oxides and hydroxides relative to the other more soluble elements. For instance Giovanoli et al. (1988) found noncrystalline alumi- num hydroxide as a product of the chemical weathering of silicates in lake sediments at a high elevation (2400 ma.s.) in the southern Swiss Alps. Additional evidence for significant chemical weathering at high elevations comes from Drever and Zobrist (1992). They studied weathering stoichiometry in a profile ranging from an elevation of 200 to 2400 m a.s.l. in the southern Swiss Alps (Canton Ticino). Accord- ing to their findings selectivity of weathering does not change with elevation, because the denudation regime ist weathering-limited (Stallard and Edmond, 1983) at all elevations due to the steep slopes. They also found no systematic elevation dependent trend in the formation of secondary minerals in the soils. For instance traces of kaolinite were found at the lowest as well as at the second highest elevation. Hence, when physical weathering becomes more intense and overrides the selective effect of chemical weathering Al and Fe can be enriched in the sediments of a receptive Weathering tracers 78 Chapter 5 lake. Similarly also Zn is often enriched in soils with respect to the related parent rock due to its low mobility and its adsorption on clay minerals, ferric Fe-oxide and organic residues (Wedepohl, 1978). Increased soil erosion thus could also lead to higher Zn concentrations in the sediments.

The Holocene record of Lej da Silvaplauna With this information temporal variations observed in the Lej da Silvaplauna elemental profiles reaching back to 10000 BP (figure 5.4) can be interpreted as follows. In Lej da Silvaplauna only K is depleted in the early and mid Holocene sediments which would indicate enhanced leaching of stable soils at that time according to Makereth (1966). Depletion of Na and Mg that should occur at the same time is not observed. However, coexisting high Ca and Sr values may support the above interpretation for they could be due to an increased precipitation of calcite in the epilimnion of Lej da Silvaplauna in response to a higher autochthonous productivity. Higher productivity in turn would be an indication of greater nutrient input and/or higher surface water temperature. Enhanced lake productivity at that time is supported by the pigment association and low C/N ratios, determined in a similar lithostratigraphic position in Lej da Segl (preliminary data). Formation of authigenic calcite thus could start because calcite, dolomite and Ca-feldspars are the first minerals to dissolve when soils develop (see above). As mentioned above considering the strong affinity of Ca for organic ligands the early and mid Holocene organic rich sediments of Lej da Silvaplauna may contain substantial amounts of Ca not associated with detrital or authigenic minerals. The profiles of Al and Fe might be taken as additional evidence for this hypothesis, since these elements are retained in the soils during weathering (cf. above). With the Neoglacial climatic deterioration (samples 2a+b) physical weathering became more intense, overriding the selective effect of chemical weathering. Thus enhanced erosion of Al and Fe enriched and Ca depleted topsoils can explain the ex- traordinarily high Al and Fe concentrations in sample 2a+b. Interpreted in the sense of Stallard and Edmond (1983) for Lej da Silvaplauna this would represent a shift from a transport-limited (non-glacial) to a weathering-limited (glacial) denudation regime. The covariation of K with Al suggests that a similar mechanism might control the K distribution. Fanning et al. (1989) and St. Arnaud and Sudom (1981) suggest that under arid conditions in response to K enrichment in the soil solution (and at the soil surface due to biocycling) re-adsorption of K by K-depleted micas or K- uptake by beidellitic smectites might occur. Although it needs further constraints this model would explain the observed covariation of K and Al. A selective K-fixa- Chapter 5 79 Weathering tracers tion in the soil could also explain why Na and Mg, do not show the same trend as K in samples 2 and 3. Only the spot-samples show a weak Na and Mg-depletion, whereas mix-samples show a slight enrichment. In contrast, under humid conditions Dean (1993) found that in the sediments of Elk Lake in particular the concentration of Na provides a good index to the degree of chemical weathering in the drainage basin. Fe and Mn cycles seem to be decoupled in Lej da Silvaplauna, because no significant correlation (r2=0.27) between them is observed in mix- or spot-samples. Instead Mn covaries with Ca (r2=0.49) and Sr (r2=0.61). In addition, Mn is enriched in the nonglacial sediments (sample 3spot) as it is in the recent sediments of the basin plain (sample PSP 90/3) (figure 5.5). This is also expressed in the Fe/Mn ratio (figure 5.4), which is very low in the nonglacial sediments and in the recent sediments of the basin plain. According to Makereth (1966) this indicates a strong differential leaching of Mn from the catchment soils. These features may also be the result of the differential mobility of Mn and Fe within the lake sediments (see below). However, severely low lake redox conditions are rather unlikely for the non-glacial early and mid Holocene, since Fe concentrations are low, in spite of presumably high sulfur concentrations - high sulfur concentrations have been observed in similar lithostratigraphic positions in Lej da Champfèr and Segl (cf. chapter 7). This indicates that redox potential was above -100 mV. Below -100 mV insoluble iron sulfides would have formed (Engstrom and Wright Jr., 1984; Jones, 1982) which would have resulted in higher Fe concentrations. Instead Fe is enriched in the Neoglacial, late

Holocene sections 2a+bmix. Fe concentrations therefore seem to be controlled by the allochthonous clastic input, whereas limnological conditions apparently only have a minor impact. However, ultimately due to the complexity of forcing mechanism Fe and Mn profiles do not yield unequivocal results about the intensity of chemical weathering of catchment soils. On the other hand, Fe concentrations correlate well with K (r2=0.66), Al (r2=0.78) and Zn (r2=0.65). This element association is most likely encountered in amphibols and biotite. In particular the unusally high Zn values in sample 2b could be due to a higher amphibol content. Wedepohl (1978) reports Zn concentrations of up to 8900 mg/kg for amphiboles in contrast to the average crustal value of 57.1 mg/kg. This interpretation is supported by the mineral assemblage of turbidites (5.3.1) that show higher amphibol contents in this section of the core (figure 5.2 and 5.8). Since these sediments were deposited during Neoglacial times with dominating input of glacigenic rockflour, Zn probably is of geogenic origin. On top of that increased soil erosion would also lead to higher Zn concentrations (see above). Weathering tracers 80 Chapter 5

Changes in Cu concentrations on the other hand are only very small. Cu therefore seems to be associated with a different source that does not change significantly during the Holocene. Since Cu-concentrations in general are very low - on average 14.1 mg/kg as opposed to a mean crustal value of 23.7 mg/kg - this can be feldspars or quartz. Almost all other minerals show higher Cu concentrations.

Complications in the interpretation of element concentrations Owing to their solubility under reducing conditions Fe and Mn may migrate within the sediment column during diagenesis. In an oligotrophic lake usually only the uppermost few centimeters of the sediment are oxidized, whereas deeper sections are under reducing conditions. This is why upward diffusion of Mn and Fe in interstitial fluids and re-precipitation in the oxidized surface layer occurs. This pumping mechanism is most effective for Mn, which is more soluble than Fe and leads to concentration maxima in the surface sediments. This mechanism probably is responsible for the Mn profiles observed in Lej da Silvaplauna and Segl. Since Fe shows no or only a weak enrichment in the surface sediments of Lej da Segl, Eh-values in interstitial waters here are probably not low enough to allow for significant reductive Fe mobilisation (Eh > approx. 200 mV). In Lej da Silvaplauna also Fe shows surface enrichment, hence Eh-values are probably below 200 mV (Jones, 1982; Stumm and Morgan, 1981). Interpreted in a different way, the observed profiles of Al, K, Ca, Na, Mg might as well be the result of the very intense leaching of rock flour in the glacier milk (Brown et al., 1996) and not in catchment soils. This leads to a depletion of the remaining particles in easily soluble species (Ca, Mg, Na) and an enrichment in Al and K (Lauriol et al., 1997) which are not readily dissolved in quickflow glacial waters (Brown et al., 1996). The contrasting K versus Na, Mg and Ca profiles could also indicate that K does not originate from the dissolution of feldspars. Instead biotite-dissolution (Fan- ning et al., 1989) or alteration of biotite to vermiculite (Drever and Hurcomb, 1986) might be the major source of K. The observed Fe profile however is inconsistent with this hypothesis because most Fe bearing minerals are easily soluble too. Consequently, Fe depletion rather than the observed enrichment would be expected for samples

2a+bmix. Alternatively the observed element concentrations could be explained by a relatively more important detrital input of dolomite, calcite or gypsum from the southeastern catchment (figure 2.3) during the early and mid Holocene. This, as a consequence of the absence of rockflour-input from glacial erosion might also lead to the observed higher Ca and Sr values. Depletion of Al and K in the same time interval Chapter 5 81 Weathering tracers might then be an effect of the dillution by carbonates. The occurrence of calcite bearing turbidites only in the early and mid Holocene sediments (c.f. 5.3.1) might argue in favour of this hypotesis. High Zn concentrations previously have been ascribed to increased soil erosion and/or higher abundance of Zn bearing minerals. On the other hand, an anthropogenically induced icrease of Zn concentrations should not be underestimated here. Peculiarly, with the beginning of neoglaciation in the late Atlanic, traces of settlements from bronze-age people can be found in the Upper Engadine (c.f. 2.8). There is also evidence for significant human-mediated deforestation at that time (c.f. 2.7). Zn concentrations in trees range from 20 to 120 mg/kg, but average Zn concentrations in the plant ash are in the range of 1000 to 1500 ppm which indicates an accumulation by more than a factor of 10 in relation to trees, sedimentary rocks and soils (Wedepohl, 1978). Therefore another possible, and maybe the most likely, source of Zn are forest clearings, especially when trees have been burned rather than cut. In summary the geochemical profiles show that although spacial and temporal differences in the element composition exist, they can not explicitly be ascribed to one single cause. This partially is a consequence of the bulk digestion procedure applied in this study. It means that changes in the relative contributions of allogenic vs. authigenic and biogenic components are not accounted for. When interpreting the element profiles in Lej da Silvaplauna Holocene sediments in terms of their implications for the catchment soil erosion and weathering history this has to be considered. Consequently, assessing the contributions of various sediment fractions becomes crucial for interpreting the lake sediment record in terms of ersional history. But even if this could be done the numerous factors that influence the retention of chemical species within the soils as well as within the lake sediments render any interpretation highly speculative.

5.3.4 Magnetic properties: Tracers for the erosion of topsoil ?

The use of magnetic properties of lake sediments as tracers in palaeoenvironmental reconstructions has been assessed in several studies. In most cases changes in bulk magnetic susceptibility were linked to variations in the allochthonous detrital input (e.g. Thompson et al., 1975). Dearing and Flower (1982) even found a positive correlation between magnetic susceptibility of material collected in a sediment trap over one year and rainfall. More detailed magnetic studies revealed that pedogenic processes can lead to the formation of finegrained particles with characteristic magnetic properties during soil developement (e.g. Walling et al., 1979). For instance Higgitt et al. (1991) were able to link changes in the magnetic properties of Petit Lac d’Annecy Weathering tracers 82 Chapter 5 sediments, to changes in land use during medieval and post-medieval times. They ascribe a decrease in the grainsize of magnetic particles to an erosion of magnetically enhanced soils in the catchment, because the magnetic changes parallel those in major element concentrations and pollen frequencies. However, it has to be considered that fine grained magnetic minerals can have other catchment sources or are produced in situ by geochemical or microbially mediated processes (Bingham Müller, 1996; Vali et al., 1987). Oldfield et al. (1985b) noted that mineral magnetic measurements can be particularly sensitive when the minerogenic content and the particle size distribution of a sediment profile changes.

Snowball (1993) suggests the S-ratio (=IRM/IRM300mT), which is a measure of the magnetic ‘hardness’ of the sediments, as a paleoclimate indicator. He studied lakes and soils in a glaciated catchment in nothern Sweden (Kårsavagge). According to his interpretation during periods of climatic amelioration (warmer and drier) the supply of minerogenic material to the lakes is reduced and sediment with a higher organic content is deposited. The slow deposition rate leads to magnetite dissolution, a low magnetic concentration and a high S-ratio. During periods of glacial advance the minerogenic sedimentation rate increases and magnetite is preserved in the sediments due to rapid burial and results in a high magnetic concentration and low S- ratio. Studies of loess profiles have shown that fine grained magnetite developes in soils under warm humid conditions (Heller et al., 1991). The autors conclude that frequency-dependent susceptibility (cfd), which is a measure for the grain size of magnetic particles (cf. section 3.4.1), seems to be a useful parameter for reconstruction of low grade weathering changes. As mentioned above, fine grained magnetite can also develop in lake sediments in situ. Bingham Müller (1996) showed that ultra- fine-grained biogenic magnetite was formed and preserved within bioturbated marls in Lake Greifen, which were sedimented under relatively low productivity conditions and permanently oxic waters. This authigenic magnetite would show up as an increase in frequency-dependent susceptibility. Since reductive dissolution occurs under anoxic conditions it mirrors the oxygen status of the hypolimnic or interstitial waters, that in turn is influenced by nutrient input and lake stratification and therefore reflects climate changes only indirectly.

The frequency-dependent susceptibility (cfd) or F-factors determined in this study in Lej da Silvaplauna Holocene sediments (figure 5.6) are compared to the glacier history as reconstructed from grain-size and varve thickness measurements per- formed on the same core by Leemann and Niessen (1994b). It is striking that with the onset of neoglaciation at 3300 BP consistently higher F-factors occur until 2650 BP. Chapter 5 83 Weathering tracers

Interpreted in the sense of Higgitt et al. (1991) this would indicate the erosion of magnetically enhanced soils. Consequently the development of stable, deep soils giving rise to the developement of fine grained magnetite in a glacier-free catchment is anticipated for earlier phases, namely the Atlantic. This also is in accordance with Dearing and Foster (1986), who found that F-factors can be used to trace the movement of topsoil. They note that naturally secondary magnetite characteristically is produced in the presence of organic matter or by burning (Walling et al., 1979) especially in free draining topsoils. Whether this erosion of topsoil is due to increasing glacial erosion or to beginning human activity (forest clearing) in the catchment (cf. 2.6) remains unclear. However, following Snowball (1993), the other possible explanation for the observed higher F-factors would be a better preservation by rapid burial of authigenic magnetite due to higher lacustrine accumulation rates as a consequence of beginning catchment glaciation. Consequently according to the author reductive dissolution of magnetite in the lake bottom waters must have occured during the non-glacial times (9400-3300 BP). Supporting evidence for the latter interpretation might be higher productivity and bottom water anoxia interpreted from pigment and sulfur data measured in a similar lithostratigraphic position in Lej da Segl (not shown). Also the fact that most of the other, although sometimes single, F-factor peaks occur in time windows were glacier advances have been identified by Leemann and Niessen (1994b) argues in favour of this interpretation. Especially the highest F-factor coincides with the most pronounced Holocene glacier advance, which occurred during the Little Ice Age around AD 1850. However, it does not fit in this picture that during the glacier advance between 2650 and 1850 BP no higher F-factors were detected. Therfore a final decision whether increases in F-factors are caused by increasing allochthonous input of magnetite from soils or by authigenesis and preservation of fine grained magnetite in the sediments themself, can not be made from these measurements. Probably an interference of both factors is recorded in the sediments. Moreover, the results of the F-factor measurements, although somewhat ambiguous, contradict the results obtained by Leemann and Niessen (1994b). From their measured SIRM-values they interprete a ‘harder’ magnetization in unit 2 in comparison to unit 3 as evidence for the absence of magnetite, but the predominant occurrence of probably detrital hematite or goethite. In contrast high F-factors in large parts of unit 2 (3300-2600, 1500-700 and 250-0 BP) suggest a significant contribution of fine grained magnetic minerals (most likely magnetite, from whatever source, authigenic or detrital) to the susceptibility signal. Additional measurements would be necessary to solve this problem. Weathering tracers 84 Chapter 5

5.4 Conclusions

Coming back to the objective of this study i.e. the assessment of the of the different measured parameters (compiled in figure 5.10) as tracers for the contribution of soil derived material - used vicariously for rainfall - to the annual sediment accumulation in Lej da Silvaplauna, the following statements can be made. None of the investigated tracers fully meets the criteria formulated in the introduction. Nevertheless, through each of them some insight is gained into the potential sediment delivery processes. The determination of magnetic properties appears to be the most promising pathway to follow. This is mainly because magnetic measurements can be carried out relatively fast and with a high spatial resolution of down to 1mm when modern techniques are used (Worm et al., 1995). This resolution is in the range of annual accumulation rates in Lej da Silvaplauna. Hence, although the measured F-factors are somewhat ambiguous (cf. 5.2.4) with the discriminating criterion used, it is likely that the detected changes show real variations in the grainsize of magnetic particles. For the use as an erosion-proxy in palaeoenvironmental reconstructions it is a prerequisite to decide whether authigenesis of magnetic minerals does occur in the lake sediments of Lej da Silvaplauna like it has been proven for the most recent eutrophic sediments of adjacent Lej da S. Murezzan by Ariztegui and Dobson (1996). The greigite they detected, however due to its coarser grainsize, would not affect the F-factor since here only ultra-finegrained (~0.03 µm) ferrimag- netic grains are detected (Oldfield et al., 1985a). In analogy to Dearing and Foster (1986) and Higgitt et al. (1991) the prove that detrital magnetite derived from the erosion of topsoil indeed is responsible for the changes in sedimentary magnetic properties, can be furnished by a multiproxy study and a more detailed magnetic analysis. Once this can be shown, measurements of frequency dependent susceptibility would be the ideal tool for the quantificaton of soil erosion rates in high resolution studies. Also other techniques that e.g. allow a mapping of even small SIRM variations with a spatial resolution of 1mm (Worm et al., 1995) offer a large potential in this respect. Even the bulk susceptibility that can now be determined in a high resolution mode might yield valuable information analogous to studies in the marine realm (Sarnthein et al., 1982; Tiedemann et al., 1994). Of course the uncertainty remains whether all soil derived material is a product of erosion by precipitation. The possibility of significant soil erosion during glacier advances can never be excluded. Ac- tually, the F-factor data shows that on centennial timescales F-factor changes often coincide with glacier highstands. This could mirror both, increased glacial erosion of Chapter 5 85 Weathering tracers ] 6 ] m yr ∗ [µ

2 12 g/cm [

MAR 0.3 0.2 0.1

Grain-size Median varves varves 48 1 [%] -factor F 02 4681012 ] kg / mg [ Zn 50 100 150 200 [%] 8.0 9.0 Al 7. 0 ] % 2.0 [ Ca 1.0 1.5 2.5 ] % [

4 Calcit e Other Stilpnomelane Serpentine Talc 02 68 2.5 3.5 1.5 Chlorite / Mica 0.5 [%] 30 50 70 90 Chlorite+Mica Chlorite+Mica 0 1000 2000 3000 4000 5000 6000 7000 8000 9000 10000 Mineral ratios, element composition and F-factor profiles in the Holocene sediments of lake Silvaplana compared to grainsize and of the same record MAR-variations Silvaplana compared in the Holocene sediments of lake F-factor profiles Mineral ratios, element composition and

925 385

BP 2775 2875 2965 3835 4620 5845 7830 8755 9415 1015 1725 Age Unit 4 3 2 1 PSP N 90/1 0 900 500 600 700 800 100 200 300 400 1000 1100 (cm) Lake Silvaplana Depth redrawn from Leemann (1993). Leemann from redrawn Figure 5.10: Figure Weathering tracers 86 Chapter 5 soils or enhanced soil erosion by increased precipitation, that of course also has an influence on glacier mass balances and causes glacier advances. Therefore, even if F- factor changes do represent soil derived material, a causal link to variations in the amount or intensity of rainfall is not feasible. The initial thought was to trace soil erosion by detecting the amount of secondary minerals formed by chemical weathering of catchment soils, especially kaolinite, in the lake sediments. Although kaolinite formation does occur in high elevation soils (Drever and Zobrist, 1992) and even was observed in a palaeosoil on the isola delta in adjacent Lej da Segl (Gehrig, 1996), this signal of soil formation apparently is not transferred to the lake sediments. In Lej da Silvaplauna kaolinite has never been detected in any of the samples. This is also true for two high elevation lakes in the southern Swiss Alps (Giovanoli et al., 1988). Neither fine-grained sections nor the clay tops of turbidites, where natural enrichment was expected did show traces of kaolinite. Therfore this mineralogical tracer has proven to be not applicable for this porpose. Aside from this, the mineralogy of turbidites proved to be a useful tool in the identification of sediment sources. In this context it yields valuable information about the response of the lake sedimentary system to changes in the extent of catchment glaciation (cf. 7.3.5) and perhaps permafrost distribution. It has been demonstrated that the turbidite system responds to changing boundary conditions with a lag time of several hunded years. Based on the turbidite record, a degradation of permafrost in the southeastern catchment is proposed for the time period 3800-6200 BP. The higher F-factors in the same time intervall could also be interpreted as reflection of this degradation. This hypothesis of course needs further investigations especially in view of the fact that a period of intense solifluction was detected at Albulapass (eastern Switzerland) by Gamper (1982) between 3300-2900 BP. Moreover, according to my knowledge, it is presently not well known what kind of material is to be expected in the lake when permafrost is degrading. It would be of great interest if this material does exhibit any characteristic features that allow a distinction from glacier or rainfall eroded material. E.g. Vonder Muehll (pers. comm.) detected a significant layer of very fine-grained sediments in a permafrost borehole on a Blockgletscher Murtel in the Upper Engadine. Geochemical sediment tracers, although potentially well suited as tracers for catchment soil erosion, are not easily determined. Especially because this study confirmed the necessity of applying a selective digestion procedure as e.g. proposed by Förstner and Wittmann (1979) or Engstrom and Wright Jr. (1984) the expenditure of time and money would be tremendous when applied in a high resolution mode. For this rea- Chapter 5 87 Weathering tracers son geochemical tracers can not realisticly be used in this manner. Moreover, any selective digestion method is operationally defined i.e. the distinction between sediment components might be blurred. Therefore, perhaps the most promising method to get a grasp on the weathering processes are single grain studies like they have been conducted by Giovanoli et al. (1988). Ideally the TEM/SEM and EDX techniques they applied to document the corrosion of e.g. feldspar and biotite grains should be combined with sequential geochemical analysis or perhaps (microprobe analysis) and XRD-techniques. A crosscalibraton with magnetic properties for a representative dataset could then yield the desired high resolution tracer. Being aware of the limited information potential of bulk sediment elemental concentrations, Ca and Zn values are probably the most promising tracers. High Zn concentrations in the time intervall 1200 to 3300 BP could represent the erosion of material derived from anthropogenic forest clearings. Intensive clearing probably by burning, of the forest is evidenced in pollen diagramms and by the occurence of charcoal particles in cores from several bogs around Maloja and St. Moritz around 3400 BP (Kleiber, 1974). This could also lead to the formation of fine grained magnetite (Oldfield et al., 1985a; Walling et al., 1979). The higher F-factors at the biginning of this time intervall might thus reflect this event as well. Moreover the sustained turbidite generation in the northwestern catchment until 2650 BP perhaps is also caused by this anthropogenic activity in the lake catchment. In this case the non- glacial turbidite signal would rather overprint the already present glacial turbidite signal than representing a real lag time in delta activity. In the Lower Engadine human impact on vegetation as evidenced by pollenanlysis dates back to as early as 5600 BP (Zoller et al., 1996). Therefore the F-factor as well as mineralogical changes in turbidite composition could as well mirror this increasing human influence on catchement erosion rates rather than a climatic influence. The surprisingly high Ca concentrations in unit 3 might offer a chance to trace the intensity of chemical erosion, since the leaching of catchment soils and rocks starts with the dissolution of carbonates. Prior to any interpretation though, it has to be determined whether the high Ca-concentrations are caused by increased allochthonous detrital input or indeed are a result of lake internal calcite precipitation. The high Ca/Mg-ratio of sample 3 points to calcite or gipsum as Ca bearing minerals rather than detrital dolomite. A decision about the origin of calcite could be achieved e.g. with SEM-techniques. If authigenic calcite does occur it also opens the possibilty of employing isotopic tracers. Stable oxygen and carbon isotopic measurements in authigenic carbonates could then yield a information about changes in the isotopic composition of the lake water. In the case of Lej da Silvaplauna, which is an oligotrophic lake with a short residence time, under glacier-free catchment condi- Weathering tracers 88 Chapter 5 tions in the early and mid Holocene, these changes should mirror the isotopic composition of the precipitation entering the lake. Owing to Lej da Silvaplauna’s geo- graphic position right on the boundary of two climate zones, where air masses from the mediterranian meet with air masses carried to the interior Alps by a northerly flow, an isotope record would offer the opportunity to reconstruct early and mid Holocene atmospheric dynamics. In conclusion for a purely chemical characterization of lake sediments the state- ment of Engstrom and Wright Jr. (1984) is still valid. They summarize that in oligotrophic lakes the primary environmental signal represented in the sediments is the composition of catchment soils, whereas in more productive lakes authochthonous production of carbon, bacterial activity, and redox variations can produce internal modification of the sediment composition that may be read from chemical stratigraphy as a history of limnological developement. Since the latter is true for some time intervalls in Lej da Silvaplauna’s history their conclusion that chemical stratigraphy has not yet reached the point where it can be used independently of other evidence to reconstruct lake history can be adopted for this study. Therefore a multi-proxy approach seems to be the adequate way to overcome these deficiency. Combining the information gained from the determination of different sediment properties will grant a better understanding of lake versus catchment driven processes. The importance of reconstructing the vegetation developement by pollen analysis especially when human impact has to be evaluated, might be underlined at this point. Chapter 6 89 Neoglacial variability

CHAPTER 6

Interannual and interdecadal variability in the Neoglacial varve record of Lej da Silvaplauna

6.1 Introduction

Together with treerings, annually laminated lacustrine sediments are the only terrestrial source of information that allows the detection of past environmental changes at a high resolution for time periods of more than 1000 years. In Lej da Silvaplauna glaciolacustrine mass accumulation rates (MAR) were determined with an annual resolution for the last 3533 years (Leemann and Niessen, 1994b). In this chapter this varve record is re-analyzed under the following aspects. 1. The analysis presented in chapter 4 demonstrated that a calibration of the Lej da Silvaplauna varve record is not easily accomplished. Although mean summer temperature can explain ~50% of MAR variability there are still 50% unexplained variance. The comparison with information available from the local archives revealed that forcing factors on sediment accumulation are diverse and variable through time. Hence a better understanding of the varve record and its significance for climatic reconstructions is needed. This topic can be approached for example by a more detailed analysis of the varve record. For this reason the sediments dated by Leemann and Niessen (1994b) were sampled again in this study in order to evaluate other potential climate proxy indicators that could be determined with an annual resolution. It is expected that accumulation rates of biogenic silica (BiSi), organic carbon (TOC) and carbonate carbon (TIC) could yield information about the productivity of the lake, which is indirectly related to the water temperature. This information may help to isolate the climatic signal encoded in the sediments. 2. It remains to be tested how this high alpine varve record compares to other high resolution records of climate (instrumental data from long term European climate stations, historical documents, treerings). In particular this comparison focuses on the time window AD 1738-1766, which is identified as an anomaly in the varve record. This should show in how far the Lej da Silvaplauna varve record is representative for a larger region (Europe) on timescales of years to decades. It can also lead Neoglacial variability 90 Chapter 6 to a better understanding for the interrelationships between different climate variables and the lake’s sedimentary record.

MAR [mg/cm2∗a] MAR 5yr RAv 0 100 200 300 400

BiSi (mg/cm2∗a) BiSi 5yr RAv 0 5 10 15

C_org (mg/cm2∗a) C_org 5yr RAv 0.0 1.0 2.0 3.0 d

C_carb (mg/cm2∗a) C_carb 5yr RAv 0.0 0.5 1.0 1.5 2.0 1994 1984 1974 1964 1954 1944 1934 1924 1914 1904 1894 1884 1874 1864 1854 1844 1834 1824 1814 1804 1794 1784 1774 1764 1754 1744 1734 1724 1714 1704 1694 1684

Figure 6.1: Profiles of a) MAR, b) biogenic Silica (BiSi), c) organic carbon (C-org) and d) carbonate carbon (C- carb) accumulation rates in the sediments of lake Silvaplana between AD 1684 and 1994. Heavy black line=smoothing with a 5 point running mean. Area shaded in light grey marks the time interval AD 1738- 1766. Chapter 6 91 Neoglacial variability

3. Because of the true annual sampling interval it is feasible to apply timeseries analysis (FFT spectral analysis) to the Neoglacial varve time-series. The advantage of spectral analysis is that it can reveal periodic forcing mechanisms on lacustrine sedimentation without a fully understanding of all the processes controlling interannual variations in MAR. This may then help to develop ideas about the forcing of MAR changes by climatic processes.

6.2 Results

6.2.1 Geochemistry and MAR of the sediments deposited since AD 1494

Biogenic silica (BiSi), total organic carbon (TOC) and carbonate contents (Ccarb) were measured on 125 samples from core PSP 94/3-0 with a resolution of 1 to 5 years. The upper part of the BiSi, TOC and MAR profiles comprising the period from AD 1864 to 1994 has already been discussed in chapter 4 (figure 4.4). Here this record is extended back to AD 1683 comprising the latest part of the Little Ice Age (figure

3.5 6.1). The measured contents are listed in appendix 6.1. For the calculation of 3.0 annual component accumulation rates 2.5 (cf. 3.4.6) turbidites were disregarded [%] 2.0 and gaps in the dataset, where no measurements were conducted, were filled 1.5 by linear interpolation between the

Organic carbon 1.0 two nearest datapoints. y = 0.1645x + 0.2724 0.5 R2 = 0.9001 BiSi and TOC profiles show a good 0.0 correlation (r2=0.90, figure 6.2). There- 0510 15 20 Biogenic silica [%] fore, because BiSi is exclusively of authigenic origin the same seems to be Figure 6.2: Biogenic silica content plotted against the content of total organic carbon in the sediments of lake the case for the majority of TOC. For Silvaplana for the years AD 1683-1994. The good correla- this reason in the time window AD tion indicates a common (autochthonous) originin of these components. The equation computed by linear regression 1710 to 1770 where no BiSi data were is used to calculate BiSi contents in the time window 1738- available, the BiSi-contents were cal- 1766 where no BiSi measurements were conducted. culated using the equation derived from the correlation with TOC-contents (figure 6.2). In addition to the BiSi, and TOC profiles the carbonate accumulation rates are shown in figure 6.1d. Prior to AD 1683 only MAR data is available from Leemann (1993). Neoglacial variability 92 Chapter 6

In figure 6.1 maxima of MAR, BiSi and Corg accumulation occur in the time intervals AD 1960-1994, 1940-1950, 1916-1925, 1884-1891, 1838-1850, 1794-1800 and 1620- 1660 (only MAR data in figure 6.4a). In contrast, the most dramatic reduction in BiSi and Corg accumulation rates and MAR is observed between AD 1738 and 1766 (shaded area in figure 6.1). This interval is investigated in more detail in this chapter. It is striking that the Ccarb accumulation rates do not show the same pattern of variation (figure 6.1d). Minimal Ccarb accumulation rates were detected between AD 1770-1794. Striking maxima occur at AD 1942-1956, 1830-1838 and 1818-1822. The correlation curve of Ccarb accumulation with MAR (figure 6.3) shows extended

11pt runcorr. MAR/C_carb 0.0 0.2 0.4 0.6 0.8 1.0 1994 1984 1974 1964 1954 1944 1934 1924 1914 1904 1894 1884 1874 1864 1854 1844 1834 1824 1814 1804 1794 1784 1774 1764 1754 1744 1734 1724 1714 1704 1694 1684 Figure 6.3: Running correlation coefficients (11 points) between MAR and the carbonate carbon content (C_carb) of lake Silvaplana sediments. Note the decreasing correlation starting at the beginning of this century (shaded area). periods with bad correlation towards younger sediments. In particular since the beginning of this century (shaded in figure 6.3) and around AD 1730 and 1770 the correlation between these two parameters gets weaker.

The turbidite record The occurrence of turbidites back to AD 1494 is shown in figure 6.4b. The largest turbidite of the core occurs at AD 1816. An unusual clustering of turbidites occurs between 1956 and 1961, whereas the longest time span were no turbidites occurred is between AD 1721 and 1771, coinciding with the time of minimal MAR (figure 6.1a) mentioned above. Turbidites that have been ascribed to historically documented flood events (Röthlisberger, 1991) or anthropogenic influence i.e. road works at the lake shore, are indicated by the year of the occurrence of the event in figure 6.4b. Chapter 6 93 Neoglacial variability

MAR (g/cm2∗yr) a 5yr RAv 0 0.1 0.2 0.3 0.4 0.5 1994 1974 1954 1934 1914 1894 1874 1854 1834 1814 1794 1774 1754 1734 1714 1694 1674 1654 1634 1614 1594 1574 1554 1534 1514 1494

Turbidites (g/cm2∗yr) b 1888 1954 1566 1793 (Landslide in Silvaplana catchment) 1845 (Road works, all lakes affected) 1868 1640 1987 1834 1820 (Road works in Silvaplana) 1496 (Landslide in adjacent valley) 1927 1771 1519 01234567 1994 1974 1954 1934 1914 1894 1874 1854 1834 1814 1794 1774 1754 1734 1714 1694 1674 1654 1634 1614 1594 1574 1554 1534 1514 1494

Figure 6.4: a) MAR record and b) the occurrence of turbidites between AD 1494 and 1994. Years were exceptionally rainfall events are historically documented are noted by the year of their occurrence in b). Heavy black in a) line=smoothing with a 5 point running mean. Shaded area marks the time interval AD 1738-1766.

6.2.2 Comparison of the MAR record with climate (proxy) data on a regional scale

The MAR data shown in figure 6.4a is a compilation of data from this study and the study of Leemann (1993). It is used for the comparison with long climate proxy records reaching back to AD 1525 from Jones and Bradley (1995); Pfister (1995), F. H. Schweingruber (pers. comm.) and Livingstone (1997) (cf. 3.8). These records contain numerous datasets most times on a seasonal basis, that were compared with the MAR record. In figures 6.5 to 6.7 only selected datasets are plotted. Subsequently, first the general correlation with respect to the MAR record is characterized for this century. Second the main focus is placed on the AD 1738-1766 MAR-minimum (shaded in figures 6.5 to 6.7).

General correlation The MAR record was compared with temperature and precipitation data from some of the longest European instrumental climate records (Jones and Bradley, 1995). Neoglacial variability 94 Chapter 6

In this study only seasonal and annual means of the records that reach back into the time period of interest (AD 1738-1766, shaded in figures 6.5 to 6.7) were considered. These are temperature records from Berlin, Stockholm, Geneva and Central England. Precipitation records from England+Wales, Hoofdorp, Milano, Padua, Lund, Marseille and Paris. The aim was to select records that are representative for maritime climate on the one hand and continental climate on the other hand. Moreover, the climate of northern Europe and of the alpine region are represented in the available climate data. For most temperature records a generally good agreement of MAR (figure 6.5a) with mean summer temperatures is observed for the time interval AD 1940-1960 (shaded in figure 6.5). The records of Geneva, Berlin and Central England (figure 6.5b-d) are shown as an example. Also a summer temperature minimum between 1910 and 1920 that correlates with the MAR is seen in most records (e.g. Berlin and Geneva, figure 6.5b+c). For the Geneva record between 1900 and 1960 a positive correlation of r2=0.64 with MAR is observed. Before 1900 for most records no longer periods of consistently good agreement with MAR were detected. In the Central England summer temperature record these features are not expressed very clearly. Instead there a better correlation with MAR in the time interval 1940-1960 is observed for the mean spring temperatures (figure 6.5e). For the precipitation records in general the agreement with MAR is not is not as good as for the temperature records. The records of summer precipitation from Hoofdorp (Netherlands) and Padua (Italy) are shown for comparison in figure 6.6b+c. The Hoofdorp record of summer precipitation (figure 6.6 b) shows a positive anomaly around AD 1890 which agrees with the MAR anomaly that has been detected for the same time. The summer precipitation record of Padua (figure 6.6c) does not show this phenomenon. MAR was compared to temperature and precipitation indices, called Pfister indices hereafter, for all of Switzerland reaching back to AD 1525 (Pfister, 1995) in figure 6.7d+e. Like for the instrumental Geneva record (figure 6.5b) the warm 1940s and the cold period around 1910 also show up in the record of summer temperature indices. Around AD 1890 slightly higher fall precipitation indices are observed which is consistent with the MAR and Hoofdorp datasets. The MAR record from Lej da Silvaplauna has been compared to several treering datasets from high Alpine settings. The records from Iffigen Alp and Suaiza are shown in figure 6.7b+c. Like these two datasets most of the treering datasets show a good positive correlation with MAR in the time window 1900-1994. Again the 1940s and 1910 anomalies are clearly detectable. The agreement of MAR with the dataset from Iffigen Alp is especially well developed (figure 6.7b). Chapter 6 95 Neoglacial variability 1994 1984 1974 1964 1954 1944 1934 1924 1914 1904 1894 1884 1874 1864 1854 1844 1834 1824 1814 1804 1794 1784 1774 1764 1754 1744 1734 1724 1714 1704 1694 1684 1674 1664 1654 1644 1634 1624 1614 1604 1594 1584 1574 1564 1554 1544 1534 1524 a

MAR (normalized) 5yr RAv -3.0 -1.0 1.0 3.0

T sum (Geneva) 5yr RAv -3.0 -1.0 1.0 3.0 c

T sum (Berlin) 5yr RAv -3.0 -2.0 -1.0 0.0 1.0 2.0 3.0 d

T sum (Central England) 5yr RAv -2.0 -1.0 0.0 1.0 2.0 e

T spring (Central England) 5yr RAv -2.0 -1.0 0.0 1.0 2.0 1994 1984 1974 1964 1954 1944 1934 1924 1914 1904 1894 1884 1874 1864 1854 1844 1834 1824 1814 1804 1794 1784 1774 1764 1754 1744 1734 1724 1714 1704 1694 1684 1674 1664 1654 1644 1634 1624 1614 1604 1594 1584 1574 1564 1554 1544 1534 1524 Figure 6.5: a) MAR of lake Silvaplana compared to instrumental records of mean summer air temperature from b) Geneva, c) Berlin, d) Central England and mean spring air temperature from Central England. Heavy black line = 5 point runing mean. Shaded intervals denote the time periods AD 1940-1960, 1910-1920, 1885-1895 and 1738-1766. Neoglacial variability 96 Chapter 6 1994 1984 1974 1964 1954 1944 1934 1924 1914 1904 1894 1884 1874 1864 1854 1844 1834 1824 1814 1804 1794 1784 1774 1764 1754 1744 1734 1724 1714 1704 1694 1684 1674 1664 1654 1644 1634 1624 1614 1604 1594 1584 1574 1564 1554 1544 1534 1524 a

MAR (normalized) 5yr RAv -3.0 -1.0 1.0 3.0

P sum (Hoofdorp) 5yr RAv -2.0 -1.0 0.0 1.0 2.0 3.0 c

P sum (Padua) 5yr RAv -2.0 0.0 2.0

Ice BreakUp St.Moritz (Julian days) 5yr RAv 110 120 130 140 150 160 1994 1984 1974 1964 1954 1944 1934 1924 1914 1904 1894 1884 1874 1864 1854 1844 1834 1824 1814 1804 1794 1784 1774 1764 1754 1744 1734 1724 1714 1704 1694 1684 1674 1664 1654 1644 1634 1624 1614 1604 1594 1584 1574 1564 1554 1544 1534 1524

Figure 6.6: a) MAR of lake Silvaplana compared to instrumental records of cumulative summer precipitation from b) Hoofdorp (Netherlands), c) Padua (Italy) and d) to the break-up dates of lake ice in adjacent lake St. Moritz. Heavy black line = 5 point runing mean. Shaded intervals denote the time periods AD 1940-1960, 1910-1920, 1885-1895 and 1738-1766.

Comparing the MAR with the dates of ice break-up (figure 6.6d) in adjacent Lej da S. Murezzan (Livingstone, 1997) shows that these two parameters are inversely correlated for the 1940s. For the 1910 cooling this correlation is weak. Chapter 6 97 Neoglacial variability 1994 1984 1974 1964 1954 1944 1934 1924 1914 1904 1894 1884 1874 1864 1854 1844 1834 1824 1814 1804 1794 1784 1774 1764 1754 1744 1734 1724 1714 1704 1694 1684 1674 1664 1654 1644 1634 1624 1614 1604 1594 1584 1574 1564 1554 1544 1534 1524 a

MAR (normalized) 5yr RAv -3.0 -1.0 1.0 3.0

Treerings Iffigenalp 5yr RAv b -150 -50 50 150

Suaiza Tessin c 5yr RAv -120 -60 0 60 120

Temp. summer (Pfister) 5yr RAv d -2.5 -1.5 -0.5 0.5 1.5 2.5

Prec. fall (Pfister) 5yr RAv e -2.5 -1.5 -0.5 0.5 1.5 2.5 1994 1984 1974 1964 1954 1944 1934 1924 1914 1904 1894 1884 1874 1864 1854 1844 1834 1824 1814 1804 1794 1784 1774 1764 1754 1744 1734 1724 1714 1704 1694 1684 1674 1664 1654 1644 1634 1624 1614 1604 1594 1584 1574 1564 1554 1544 1534 1524

Figure 6.7: a) MAR of lake Silvaplana compared to proxy climate records from b) late wood densities of treerings at Iffigen Alp, c) late wood densities of tree-rings at Suaiza (Ticino), d) mean summer air temperature indicees and e) cumulative fall precipitation indicees from Pfister (1995). Heavy black line = 5 point runing mean. Shaded intervals denote the time periods AD 1940-1960, 1910-1920, 1885-1895 and 1738-1766. Neoglacial variability 98 Chapter 6

Correlation in the time window AD 1738-1766 The only temperature records that cover the time period of interest for this study (1738-1766, shaded in figure 6.5) completely are the Central England and the Berlin record (figure 6.5c-e). The Berlin record is discontinuous but seems to show a large shift from the lowest to the highest summer temperatures on record in this time window (figure 6.5c). In contrast the Central England summer temperatures (figure 6.5d) do not show an anomaly during this time. Instead the spring temperatures (figure 6.5e) show a minimum in the first two thirds of this time window and a maximum towards the end. The only two precipitation records that reach back to AD 1738 are Hoofdorp (Neth- erlands) and Padua (Italy) (figure 6.6b+c). For the Padua record of summer precipitation a minimum occurs in the initial part of the AD 1738-1766 time interval. For the rest of the interval no anomalous values are observed in this record. Similarly the Hoofdorp record also shows an initial minimum. But in this record this is followed by above average values of summer precipitation. For the Pfister indices (figure 6.7d-e) in the time period between AD 1738 and 1766 the lowest fall precipitation indices of the entire record are observed. In contrast summer temperatures do not show an anomaly in this time window. Instead fall and spring temperatures (not shown) exhibit very low values in this time interval too. In the treering datasets a striking exception to the usually weak correlation with MAR prior to 1900 is the good positive correlation between MAR and late-wood density of treerings at Iffigen Alp and Suaiza (Ticino) (figure 6.7b+c). All other treering datasets do not show a significant correlation with the MAR record during this time interval. The Suaiza record is the only one where values are below the average for the entire AD 1738-1766 period.

6.2.3 Time series analysis of the Neoglacial varve record

A spectral analysis (cf. 3.8) was done for the MAR dataset comprising the last 3533 years and the discontinuous turbidite record in the same time interval. MAR, Corg, BiSi and Ccarb datasets for the period AD 1683-1994 were also analyzed. The most pronounced peaks that were detected in the power spectra are listed in table 6.1 sorted for the above mentioned datasets. Relevant power spectra are compiled in figure 6.8. The spectral analysis of the 3533 yr long MAR record revealed no dominant frequencies above the 90% confidence level during the Neoglacial section of the varve Chapter 6 99 Neoglacial variability record. The frequency at 210 years is below the significance level. But in the discontinuous turbidite record peaks with a spectral power above the confidence level are detected at frequencies of 217, 247, 335, 411, 463 and 630 years (figure 6.8a, table 6.1). However, these peaks are rather broad.

120 45 Turbidite (0-3550 BP) C_org (AD 1683-1994) a accumulation rates 40 c 100 accumulation rates 35

80 30

25 60 Power

Power 20

40 15

10 20 5

0 0

0 200 40 0 600 800 1,000 0 50 100 150 200 250 300 Period [yrs] Period [yrs]

20.0 14 b d C_carb (AD 1683-1994) 17.5 accumulation rates 12 15.0 10 Error prob. = 0.001 12.5 Error prob. = 0.005 Error prob. = 0.01 8 10.0 Error prob. = 0.1 Power Power 6 7.5 Error prob. = 0.5

4 5.0

2 MAR (AD 1683-1994) 2.5

0 0.0 0 50 100 150 200 250 0 50 100 150 200 250 300 Period [yrs] Period [yrs]

Figure 6.8: Lomb transformation of a) the discontinous turbidite record from lake Silvaplana between 0 and 3534 BP and lake Silvaplana records in the time window AD 1683-1994 of b) MAR, c) organic carbon accumulation rates and d) carbonate carbon accumulation rates. Periodicities of detected spectral peaks are compiled in table 6.1.

In the AD 1683-1994 window the MAR data (figure 6.8b) shows spectral peaks above the 90% confidence level at 75 and 53 years. Peaks below the confidence level occur at 15, 25 and 111 years. Below the confidence level all of the AD 1683-1994 datasets, show spectral peaks at 48-53 years (table 6.1). In addition, depending on the dataset, spectral power is concentrated around 100, 70, 51, 31, 24 and 15 years (figure 6.8a-d, table 6.1). All of these peaks are considered here, although they are

Table 6.1: Spectral peaks detected in the Lomb transformation of different time series from lake Silvaplana. Spectral peaks centered at Dataset >250 150-250 ~100 ~70 ~51 ~31 ~24 ~15 Corg accumulation (1683-1994 AD) 92 65 51 24 15 BiSi accumulation (1683-1994 AD) 100 64 50 15 Ccarb accumulation (1683-1994 AD) 111 61 48 31 15-18 MAR (1683-1994 AD) 111 76 53 25-28 16 MAR (0-3533 BP) 210 Turbidites (0-3533 BP) 630/463/411/335 247/217/204 Neoglacial variability 100 Chapter 6 sometimes below the significance level (Anderson and Koopmans, 1963). It is pointed out here that although periodicities are given as single numbers in years, each of these numbers rather represents a band of possible periodicities. Depending on the method used for frequency analysis the resulting periodicities might be slightly different. In order to separate high frequency from intermediate and low frequency signals the long (3533 yrs) MAR dataset was dispersed into 13 different frequency bands, applying a Gaussian bandpass filter (figure 6.9). The resulting separation into different frequencies reveals pronounced signal modulation that changes with the frequency band considered and with time. Regarding the bandpass filtered record in terms of variability through time, reveals that temporal amplitude variations occur in different time windows for different frequency bands. The highest amplitudes especially for the higher frequencies (5-63 yrs) occur between AD 1000 and 1994 (shaded in figure 6.9). A second interval with higher amplitudes in the high frequency bands is observed between 600 BC and AD 100 (shaded in figure 6.9). For the time interval before 600 BC there are only very small amplitudes at all frequencies. Only the 51 year frequency band shows slightly higher amplitudes between 1250 and 1500 BC. Considering the low frequency bands (63-990 yrs) a slightly different picture emerges. Highest amplitudes in the 100, 150 and 207 year frequency bands occur in the time windows AD 1100-1600 , AD 1300-1800 and AD 1600-1990, respectively. A similar succession, i.e. initially pronounced 100 yr cycles give rise to 150 yr cycles and finally 207 yr cycles, also occurs in the time window 600 BC to AD 400. Outside of these time windows only minor amplitude variations are observed. It is noted that around AD 1600 the most obvious change occurs. There the 100 yr and slightly earlier the 51 yr periodicities stop abruptly, whereas the 24 yr and 207 yr periodicities suddenly occur with the highest amplitudes on record (areas shaded in dark grey in figure 6.9).

6.3 Discussion

6.3.1 Evaluation of geochemical sediment parameters as climate proxies

Persisting high MAR values (figure 6.4a) between AD 1794 and 1874 and from AD 1620 to 1660 mark the last two maximum glacier extensions during the LIA cold period (Zumbühl and Messerli, 1980). In order to obtain additional information about Chapter 6 101 Neoglacial variability the impact of these cold intervals for instance on lake water temperatures and productivity, geochemical sediment properties were determined (figure 6.1).

-0.15 5 yr (3-11) 0 11 yr (8-18) 0.15 15 yr (13-19) 0.3 24 yr (18-36) 0.45 31 yr (25-42) 0.6 43 yr (38-50) 0.75 51 yr (40-70) 0.9 63 yr (48-95) 1.05 100 yr (80-130) 1.2

150 yr (130-180) 1.35

207 yr (175-260) 1.5 350 yr (290-420) 1.65 512 yr (340-990) 1.8

1.95

Year -6 994 794 594 394 194 -206 -406 -606 -806

AD/BC 1994 1794 1594 1394 1194 -1006 -1206 -1406

Figure 6.9: Gaussian bandbass filtering of the neoglacial varve (MAR) record from Leemann (1993). Intervals shaded in light grey mark time periods with more pronounced periodicity at AD 1994-1000 and AD 100-600 BC. Areas shaded in dark grey indicate a switch in periodicities occurring around AD 1600. See text for explanation. Neoglacial variability 102 Chapter 6

Initially it was assumed that BiSi and Corg accumulation rates (figure 6.1b+c) could serve as tracers for the lake productivity. Especially BiSi which is enriched by a factor of up to 8 in the recent sediments with respect to the pre-eutrophic period (cf. chapter 4), was expected to show reduced values during colder periods, like the last part of the LIA (AD 1800-1850). This assumption is based on the fact that lake productivity is, amongst other factors, a function of water temperature. Photosyn- thetic activity of algae in the epilimnion increases with the availability of light until it reaches an optimum. This photosynthetic light optimum decreases with decreasing water temperature (Schwoerbel, 1980), e.g. by a factor of 2.6 between 17 and 5°C for the diatom Asterionella (Talling, 1957 in Schwoerbel, 1980). Therefore, primary production rates are directly related to the water temperatures in the epilimnion. However, although BiSi concentrations are lower in the last cold interval of the LIA between AD 1800-1880 (appendix 6.1), no reduction is observed when annual BiSi accumulation rates are calculated (figure 6.1b). The same is true for the Corg accumulation rates (figure 6.1c). Instead the lowest accumulation rates for these components are observed between AD 1738 and 1766. Hence, the idea of using BiSi and Corg accumulation rates to trace temporal changes in productivity and relate them to climate has to be abandoned. This is so mainly because the large increase in the accumulation rates that occurred since 1960 (figure 4.4) shows that the availability of nutrients and not water temperature is the major factor controlling primary production in Lej da Silvaplauna. Therefore, before 1960 when the lake system is oligotrophic, Corg and BiSi accumulation is very low and variations that might have been caused by changes in water temperature are not detectable. Any significant variation in these components hence is ascribed to changes in anthropogenic nutrient input.

Ccarb accumulation rates as glacier size proxy The fact that a poor correlation between bulk MAR and Ccarb accumulation rates is observed for some time intervals, especially after 1900 (figure 6.3), indicates that Ccarb accumulation probably integrates a different climatic signal than MAR. Refer- ring to the calibration dataset presented in chapter 4, reveals that the correlation of mean summer air temperature in Segl-Maria with Ccarb accumulation rates is actually more persistent than with bulk MAR (figure 6.10). In particular Ccarb exhibits consistent positive correlation with mean summer temperature from 1900 until 1960, whereas for MAR this correlation decreases between 1922 and 1938 (cf. chapter 4). Therefore, Ccarb might even be a better proxy-indicator for temperature controlled glacier size variations than MAR. This probably because the occurrence of carbonates is restricted to a smaller area in the catchment of Lej da Silvaplauna (figure 7.14). Chapter 6 103 Neoglacial variability

MAR / T_sum (Sils) 21pt runcorr C_carb / T_sum (Sils) 21pt runcorr -1 -0.5 0 0.5 1 1994 1984 1974 1964 1954 1944 1934 1924 1914 1904 1894 1884 1874 1864 Figure 6.10: Running correlation coefficients (21 points) of the mean summer temperature at Segl-Maria with MAR (thin line) and with carbonate carbon accumulation rates (heavy line).

It therefore might be less influenced by changes in summer rainfall and more by glacial erosion. However, it has to be recalled here that Ccarb values were not measured with annual resolution for the entire record. Especially prior to 1900 values result from interpolation and mixed samples of 2 to 5 years. Consequently the Ccarb record in this section of the core is strongly influenced by MAR which have been determined with annual resolution. Therefore, in order to confirm the above hypothesis the determination of Ccarb accumulation rates with true annual resolution is necessary. The applicability of Ccarb accumulation as summer temperature proxy remains to be tested.

6.3.2 Event stratigraphy and constraints for a AD 1738-1766 anomaly Event stratigraphy In order to constrain the varve chronology of Leemann (1993) by an event stratigraphy the turbidite record was compared to the record of historically documented flood events from Röthlisberger (1991) and other activities in the lake catchment that are relevant for sediment accumulation. This was done especially in view of the uncertain origin of the anomaly between AD 1738 and 1766 (figure 6.1a-c). Figure 6.4b shows that for each historically documented event in the time window AD 1494 and 1994 an event layer is found in the sediment record. This however, does not mean that every turbidite that is observed in the sediment section relates to a flood event (cf. Siegenthaler and Sturm, 1991). Flood events are either caused by heavy rainfall or by intensive snow and/or glacier melting. As can be seen in figure 6.4b there are by far more turbidites than documented flood events. The same was observed in the sediments deposited during the last 1000 years in Lake Uri by Siegenthaler and Sturm (1991). Moreover, only a single spot in the basin is consid- Neoglacial variability 104 Chapter 6 ered. Hence, the turbidite thickness, given as accumulation rates in figure 6.4b, is unrelated to the severity of the flood event. Siegenthaler and Sturm (1991) state that an exceptionally high suspension load is needed to produce flood turbidites that are able to carry coarse grained material into the basin. In fact the largest historically documented flood occurred in AD 1566 when e.g. the mineral springs of St. Moritz were completely destroyed (cf. 2.7). This event ‘only’ left behind a relatively thin event layer in the sedimentary record, as did the last centennial flood in 1987 (cf. 2.5). This, on the one hand, supports the assumption that turbidites often are triggered by hydrological extreme conditions. Therefore, when they occur at the ‘right place’ in the stratigraphic record they can serve as additional constraints for the proposed varve chronology. On the other hand, it demonstrates the importance of delta dynamics (e.g. retention of sediment during floods and releaseas a turbidity curr ent with a lag time when a threshold value of sediment load is reached), which makes turbidite generation highly unpredictable.

The time period AD 1738 - 1766 In the comparison with climate proxy data the focus is placed on the time interval of unusually low MAR, Corg and BiSi accumulation rates between AD 1738 and 1766 (figure 6.1a-c). This time interval is sandwiched between two cold phases of the LIA that occurred in the first half of the 17th and 19th century (e.g. Pfister, 1980). These two cold phases caused glacier advances (e.g. Zumbühl and Messerli, 1980) and therefore are also clearly recognized in the sediment record as sections of persis- tently high MAR values (figure 6.4a). The minimum accumulation rates between these two glacier advances at AD 1738-1766 are especially puzzling because a link to meteorological extreme conditions is not immediately obvious. These tree decades have never been highlighted in other climate proxy studies as being extraordinary in any respect. Therefore, first it was tested whether the unusually low MAR in this time interval could have other than climatic reasons. A possible cause for such a sedimentary feature is e.g. human activity in the lake catchment. For instance damming or deviation of the major tributary would also result in reduced sediment input and hence lower accumulation rates. Misinterpretation of the sediment record is another possibility. For instance the respective section of low MAR could also be a turbidite or a homogenite. There are several arguments against these possibilities. (1) A misinterpretation seems unlikely because an X-ray radiograph of the respective sediment section confirms the very fine lamination that was counted by Leemann (1993). (2) Grainsize analysis shows a slight reduction in grainsize median and no gradation Chapter 6 105 Neoglacial variability which precludes the interpretation as a turbidite. (3) Ccarb accumulation rates do not change in this interval (figure 6.1d), indicating that the sediment source did not change either. (4) In the available historical documents no evidence for a modification of the sediment supply by human activity in the catchment was found. (5) The turbidite record matches the record of historically documented floods also before AD 1738 (figure 6.4b) and thus supports the varve chronology of Leemann (1993). Based on the presented evidence it is proposed that low MAR between AD 1738 and 1766 were forced by a change in climate. How this compares to other climate archives is one of the main topics of the following sections.

6.3.3 Regional significance of the MAR climate record with special consideration of the AD ‘1738-1766 anomaly’

Considering the results presented in chapter 4, it is not to be expected that the MAR record shows a consistent correlation with any of the long-term climate records presented above for a longer period of time. This is simply because in chapter 4 it was shown that the MAR signal is not solely controlled by a single climate factor. Nevertheless, it was demonstrated that for meteorological extreme conditions like the warm 1940s this seems to be the case (figure 4.6). Hence it is perhaps most reveal- ing to focus a comparison with European climate data on these extreme time intervals that show up as such in the MAR record. The time interval AD 1738-1766 is such an anomaly.

Regional significance of the 1940s anomaly Together with the MAR record, all the other climate proxy records from Switzer- land, i.e. temperature indices (Pfister, 1995), treering densities (F.H. Schweingruber pers. comm.) and lake ice break-up dates (Livingstone, 1997) do indicate unusual warmth for the 1940s very clearly (figures 6.6d and 6.7). In the other European temperature records (figure 6.5c-e) this phenomenon is not so pronounced but still detectable. Therefore, this seems to reflect an amplification of this global signal (Jones and Wigley, 1990) in the Alpine region as suggested by Beniston et al. (1994). In general the time period 1940-1960 shows a positive correlation of summer and/or spring temperatures with MAR because during this 20 year period a pronounced warming/cooling cycle occurred (Beniston et al., 1994) which is detectable in most European temperature records. Based on these observations it is assumed that meteorological extreme conditions persisting for more than a couple of years occur simultaneously all over Europe and Neoglacial variability 106 Chapter 6 thus are reflecting mesoscale climatic anomalies. How (e.g. in which season) these anomalies are expressed locally in instrumental and proxy climate records differs spatially and with the proxy indicator considered (cf. figure 6.5e). In the Lej da Silvaplauna MAR record these anomalies are represented as pronounced peaks. Apart from the warm 1940s other comparable MAR anomalies occur e.g. around AD 1910 and 1890 (figure 6.5a). For the AD 1890 anomaly it has been suggested in chapter 4 that it represents the response to increased summer and fall precipitation (figure 4.6). This is constrained by increased fall precipitation indices encountered in the Pfister record and increased summer precipitation at Hoofdorp at this time (figure 6.6b and 6.7e). Based on the available data all maxima and minima that are observed in the MAR record could be ascribed to either temperature or precipitation anomalies or both. Consequently, given the above assumption of mesoscale coincidence, European climate records should also show an anomaly between AD 1738 and 1766 were the Lej da Silvaplauna MAR record exhibits a pronounced minimum.

Regional significance of the AD 1738-1766 anomaly It was suggested above that changes in sedimentary BiSi and Corg accumulation rates mirror variations of the nutrient input into Lej da Silvaplauna. Nutrient input in turn most likely is linked to rainfall in summer and fall through the rate of soil erosion. This hypothesis hence would explain high BiSi and Corg accumulation around 1890 (cf. chapter 4) as a response to higher summer rainfall. Following this hypothesis the lowest nutrient input on record must have occurred between AD 1738 and 1766. For this time interval the Pfister fall precipitation indices show the lowest values of the entire record (figure 6.7e). The MAR record from Lej da Silvaplauna was compared to treering datasets (1) because treering data are available from comparable high Alpine environments, and (2) late-wood densities of trees growing at the alpine timberline are mainly influenced by the mean summer air temperature (Schweingruber, 1987b). It was shown in chapter 4 that this climate factor also significantly influences lacustrine MAR. Therefore the good agreement between low MAR and reduced late wood densities at Iffigen Alp and Suaiza (figure 6.7b+d) in the time window indicate lower summer temperatures for both locations. The fact that the European instrumental climate records that reach back to this time period also show anomalies in this time window is viewed as evidence for the occurrence of an unusual climate during this time not only in the Alpine region but also in other parts of Europe. Apparently climate anomalies are expressed differ- Chapter 6 107 Neoglacial variability ently (e.g. in a different season) in other parts of Europe (e.g. cold springs in the Central England record, large shift in summer temperatures in the Berlin record, figure 6.5c+e). From the evidence presented above the following interpretation can be given for the climate of the Upper Engadine in the time window AD 1738-1766. Very low MAR during this time can be the result of both, low summer temperatures and reduced precipitation (rainfall) during late summer and fall. Both parameters have a major impact on MAR variations on an annual scale as has been demonstrated in chapter 4. Therefore their concomitant minima are likely to result in very low MAR. The good agreement with two high alpine treering records, especially the one from Suaiza which is fairly close to Silvaplauna, suggests that low summer temperatures might have been the major reason for the observed reduction in MAR. It seems that the occurrence of low summer temperatures was restricted to the Alpine realm or even only to the southern and the high Alps influenced by southern flow. This is proposed because the Pfister temperature indices, representing mean temperatures av- eraged over all of Switzerland, do not show a significant cooling of summer temperatures in this time interval. Additional evidence for a cooling ‘event’ in the high Alps comes from a short glacier advance of Grindelwald Glacier between AD 1769 and 1778 (Messerli et al., 1978). Accounting for the lag time of 20 years for this large glacier, this advance could be the response to the above mentioned period of summer cooling. Considering the model for glacier movement that is presented in chapter 7, it seems also possible that the MAR reduction between AD 1738 and 1766, i.e. in the middle of the LIA, represents a period of reduced basal shear stress and thus reduced erosion capacity of the LIA glaciers. In chapter 7 it is proposed for the YD glaciers that reduced MAR mirrors low basal shear stress, and thus is indicative for cold and dry continental climate. For the LIA Pfister (1980) and Röthlisberger (1991) report continental tendencies (cool, dry winters and warm, wet summers) in the Swiss climate between AD 1731 and 1811 based on historical reports about meteorological extreme conditions. Together with the evidence presented above it is proposed that the AD 1738-1766 MAR minimum represents cold, dry continental climate conditions with glaciers having a low mass balance gradient.

MAR and lake ice break-up dates The timing of ice break-up on all three Upper Engadine lakes is strongly related to local and regional surface air temperatures centred on the middle of April and integrated over 4-8 weeks (Livingstone, 1997). The inverse correlation between ice break- Neoglacial variability 108 Chapter 6 up dates and MAR (figure 6.6d) found for the warm 1940s is explicable, since an early break-up should result in an prolongation of the time period were sedimentation is possible and thus lead to higher MAR. The MAR minimum around 1910 which relates to a summer temperature cooling is not reflected as clearly in the ice break-up dates. Perhaps cooling took place in the second half of the summer and thus did not affect break-up dates that much. The MAR peak around 1890 is not reflected in the ice break-up dates, which constrains that this peak is caused by increased summer precipitation and thus is unrelated to temperature.

6.3.4 Forcing mechanisms on varve deposition and the possibility of a non-linear response of varve formation Non-linear response of the MAR to climatic forcing ? The spectral analysis of the 3533 yr long MAR record revealed a very noisy signal at high frequencies which is rather characteristic for a white noise spectrum. This, however does not mean that varve deposition necessarily is completely random. To me it rather indicates that varve generation is a highly complex process, controlled by several interacting influences and not simply responding linearly to one single forcing factor. This has already been raised in chapter 4. The varve record therefore should show the characteristics of non-linear dynamic systems (Smith, 1994). A first indication for a non-linear response of the depositional system to the climatic forcing factors, is the fact that MAR values do not show a Gaussian distribution (King, 1996), figure 6.11a). Compared to a Gaussian distribution with the same mean and standard deviation, also shown in figure 6.11a, the MAR data is preferen- tially spread towards higher values (positive skew of 1.56) and strongly leptokurtic i.e. more peaked (kurtosis of 3.16). As a result, median and mode are shifted towards lower MAR values with respect to the mean.

180 10000

a MAR 0-3534 BP b 160 (7 yr RAv) Gaussian distribution 140 (7 yr RAv) 1000 120

100 100 80 Frequency 60

Cumulative frequency 10 40

0 1

0 10 100 1000 50 -50 100 150 200 250 300 350 400 -100 MAR [mg/cm2∗yr] MAR [mg/cm2∗yr]

Figure 6.11: a) Frequency distribution of the measured MAR values (heavy line) for the last 3534 years with a synthetic gaussian distribution (thin line) with the same mean value. b) double log plot of MAR versus cumulative frequency. Both data-analysis point to a non-linear behavior of the depositional system. Chapter 6 109 Neoglacial variability

Another indication for non-linearity arises when the cumulative frequencies of the varve record are plotted against MAR. In log-log space the resulting curve ap- proaches a straight line over a certain range and is curved beyond this range (figure 6.11b). Approximation by a power law yields a correlation coefficient of r2=0.78. The power law hence only describes the straight part of the curve whereas the curved part might be described by an inverse power law. This could be taken as an indication of the chaotic nature of the depositional system (Çambel, 1993). The appropriate way to analyze such systems is to use the analytical methods of higher order statis- tics (King, 1996) and chaos theory (e.g. Çambel, 1993), which is beyond the scope of this study. Therefore, the possibly non-linear response of the system to climatic forcing complicates the determination of transfer functions as has already been presumed in chapter 4.

Intervals of more pronounced periodicity By Gaussian bandpass filtering and the dynamic frequency analysis presented in Leemann (1993) it was discovered that in the time intervals AD 1000-1994 (0-950 BP) and 600 BC - AD 100 (1850-2550 BP) high and intermediate (5-63 yr) frequencies are more pronounced (shaded areas in figure 6.9). These are precisely the time intervals were thicker varves are observed i.e. glacier advances occurred (Leemann and Niessen, 1994b). It is therefore suggested that increased glacier cover in the lake catchment works as an amplifier for periodicity. Although periodic forcing is always there, it is only detectable when the sedimentary contrasts caused by this forcing are higher than the background noise. The difference between the amount of suspension (controlling MAR) released in a warm summer with respect to a cool summer is greater for a large glacier compared to a small one. Apparently for Lej da Silvaplauna the threshold above which periodicities are detectable was only reached in the two time windows with larger glaciers mentioned above. This would make sense at least for periodicities at 11, 20-25, 40-50 and 200-220 yrs that have also been detected in other sedimentary archives (Anderson, 1993b; Livingstone and Hajdas, 1998 in press) and are believed to be related to variations in solar irradiance (Anderson, 1993b; Attolini et al., 1988; Stuiver and Braziunas, 1993). In addition, many meteorological data sets, e.g. the Central England temperature record show periodicities around 15 years and at 30-40 years (Ghil and Vautard, 1991; Plaut et al., 1995). Considering this, it is not surprising that during times of higher catchment glaciation the modulating effect of periodic variations in solar irradiance leaves an imprint in the glaciolacustrine sediment record, since solar irra- Neoglacial variability 110 Chapter 6 diance is one of the main factors controlling the rate of glacier melting (Paterson, 1994). However, whether changes in solar irradiance really are large enough to produce significant climate changes in the earth’s atmosphere is still debated (Covey and Hoffert, 1997). For instance Reid (1997) suggests, that solar forcing and anthropogenic greenhouse-gas forcing made roughly equal contributions to the rise in global temperature that took place between 1900 and 1955. Mann et. al (1998) identified changes in solar irradiance as the dominant cause of climate variability between AD 1600 and 1900, whereas greenhouse gases emerge as the dominant forcing during the twentieth century. It can be assumed that during time periods with smaller catchment glaciation, where no periodicities are detectable in the varve record, randomness is introduced by the prevalence of forcing factors other than solar irradiance and the temperature changes related to this. Such forcing factors would be e.g. rainfall, glacier dynamics, upstream glaciofluvial processes, etc.. In addition, it was uncovered by bandpass filtering that an abrupt change in the style of varve deposition took place around AD 1600 in Lej da Silvaplauna (windows shaded in dark grey in figure 6.9). For exactly the same time, Biondi et al. (1997) detected an abrupt shift to interdecadal oscillatory modes (12 and 25 yr) in the Santa Barbara Basin (SBB) varves. The authors mention two global physical mechanisms that could be responsible for this change. One is the oscillation of the sub- tropical Pacific gyre and the other one is the cyclicity of solar irradiance changes. Biondi et al. (1997) also state that interaction between these two (or even more) factors might produce the observed pattern, for interaction can result in either dampen- ing or enhancement of the decadal-scale oscillations. Because of this striking syn- chronism of the abrupt change at AD 1600 in Lej da Silvaplauna and SBB varves, it is tempting to propose that both records respond to the same hemispherical (?) atmospheric forcing.

Time series analysis The long MAR record did not show any significant periodicities (figure 6.8) as already stated by Leemann (1993). In contrast, the spectral analysis of the short term records of MAR, Corg, Ccarb and BiSi accumulation for the last 312 years (AD 1683- 1994) yield several climatically relevant periodicities (figure 6.8, table 6.1). Possible reasons for these periodicites are given below. (1) Spectral peaks at 207 (turbidite record), 40-50 and 25 years that have been related to variations in solar irradiance by other authors (e.g. Anderson, 1993b; Stuiver and Braziunas, 1993) were detected in the turbidite record and the AD 1683-1994 Chapter 6 111 Neoglacial variability

MAR record. The peaks at 207 (only turbidites) and 25 years do not show a particularly strong spectral power. This indicates that although solar irradiance during certain times certainly exerts an impact on glaciolacustrine sedimentation because it controls the rate of glacier melting (Paterson, 1994), it must be considered that other factors can interfere with and/or overprint solar forcing. Typically solar cycles, i.e. the 88-year ‘Gleissberg’ cycle, the 22-year ‘Hale’ cycle and the 11-year ‘Schwabe’ cycle (Attolini et al., 1988), are not clearly developed in most of the datasets analyzed. However, it is presently not clear, how important solar variability is for changes in the earth’s atmosphere (Covey and Hoffert, 1997). Instead periodicities that were ascribed to modulation of solar forcing by changes in the North Atlantic thermohaline circulation, i.e. 13-16 years and ~50 years often occur in the datasets (table 6.1). (2) A significant peak that occurs in the AD 1683-1994 MAR and TOC datasets (figure 6.8) is centred at 52 years. Output from a coupled ocean-atmosphere model suggests that irregular oscillations of the North Atlantic thermohaline circulation occur with a periodicity of around 50 years, but are observed in a fairly broad band around this time scale (Delworth et al., 1993). Also Stuiver and Braziunas (1993) detected a 51 year periodicity in a tree-ring derived ∆14C record. Atmospheric ∆14C changes, besides being the result of solar and geomagnetic forcing are also caused by ocean perturbations. Particularly the intensity of the North Atlantic thermohaline circulation and its effect on the rates of upwelling seem to be important players in the global atmospheric ∆14C budget (Stuiver and Braziunas, 1993). Therefore changes in ocean circulation, perhaps triggered by changes in solar irradiance, seem to be the most likely explanation for the 51 year periodicity. How this pattern, through changes in atmospheric circulation, is translated into the sediments of Lej da Silvaplauna remains a matter of speculation. (3) For the significant peak around 70 years in the AD 1683-1994 MAR and TOC datasets (figure 6.8), an explanation can not be given at the moment. (4) Peaks below the significance level occur at periodicities of 13-15 and 20-25 years (figure 6.8). These periodicities are also often found in time series of meteorological data. Ghil and Vautard (1991) and Plaut et al. (1995) identified interannual and interdecadal climatic oscillations with period length of 5-6, 7-8, 15-16 and 20-25 years in the global air temperature time series (Jones and Wigley, 1990) and the Cen- tral England Temperatures (Manley, 1974). Interannual oscillations are probably linked to global aspects of the ENSO phenomenon (Ghil and Vautard, 1991), whereas interdecadal periodicity (15-16 and 20-25 yrs) is probably also attributable to changes in the thermohaline circulation or the wind-driven circulation (Delworth et al., 1993). Neoglacial variability 112 Chapter 6

(5) A periodicity of 31 years above the 90% confidence level was detected in the AD 1683-1994 Ccarb record. This is in the same range than the periodicity Renberg et al. (1984) detected in varved sediments of Lake Judesjön in northern Sweden (30-40 years). The authors suggest that it is linked to the weather situation during summer in northern Sweden, which fluctuated between being sunnier/warmer and cloudier/ colder during the past to millennia (still does ?) with a period length of 30-40 years. Because periodicities in this range also occur in several time series of meteorological data e.g. Central England (36 yrs) and Geneva (40 yrs), a link to large scale atmospheric circulation is proposed. (6) Spectral peaks detected for the turbidite are not very sharp. Because these peaks have never been found in other sedimentary records and no explanation for their occurrence is apparent they are just mentioned here without further discussion. Only the peak at 217 yr together with one at 204 yr that is barely significant (figure 6.8), are close to the 207 yr periodicity (see above) that is believed to be associated to solar modulation perhaps modified by ocean circulation (Anderson, 1993b; Stuiver and Braziunas, 1993).

6.4 Conclusions

The detailed analysis of the Neoglacial varve record from Lej da Silvaplauna first presented in Leemann (1993) rises several new interesting points and highlights possible future directions of research. • Among the geochemical tracers BiSi and Corg accumulation rates apparently respond to changes in the nutrient input that are not necessarily linked to climate changes. No response of these parameters to the LIA cooling was detected. Differ- ent from that, Ccarb accumulation rates exhibit a pattern of variation that seems to respond to changes in summer temperatures even better than bulk MAR. To elucidate the potential of Ccarb accumulation rates as a climate proxy indicator it is proposed to determine this parameter in a high (annual) resolution mode especially prior to 1900. • Meteorological extreme conditions are reflected especially well in the MAR record of this century as already has been suggested in chapter 4. They are correlable on a regional scale with instrumental climate records and with treering, historical records from the Alpine region. When other European climate records are considered, this correlation gets weaker, but is still present. Anomalies neither occur always in the same season nor are always reflected in the same parameter. Chapter 6 113 Neoglacial variability

• Between AD 1738 and 1766 a period of unusually low MAR has been identified in the varve record, which is constrained by the event stratigraphy. A close agreement with low latewood densities at Iffigen Alp and Suaiza is observed. This lends support to the idea that this feature of the MAR record represents a reduction of summer temperatures. The comparison with other European long term records of climate suggests that low temperatures probably were restricted to the southern and high Alps and were accompanied by low fall precipitation for Switzerland. It is suggested that this interval represents a mid LIA period of reduced glacier basal shear stress in response to a cold, arid continental climate. In order to obtain a better picture of the nature and the regional significance of this anomaly it is necessary to track it in other archives. In particular, other treering series and high resolution stable isotope stratigraphies from lake carbonates from the Swiss plateau would be of interest in the first place. • The spectral analysis revealed that periodicities are not strongly developed in the Lej da Silvaplauna Neoglacial varve record. Only in certain time intervals weak periodicities are detectable. The occurrence of periodicities at 207, 52 and 25 years points to a link with solar irradiance changes. The occurrence of periodicities at 31 and 13-15 yrs demonstrates that also other forcing factors certainly exert an influence on varve deposition, through their interference with solar cycles. One possible modulating factor is the strength of North Atlantic thermohaline circulation. • Bandpass filtering and frequency analysis of the MAR data revealed that around AD 1600 an abrupt change in the dominant oscillatory modes from ~100 year cycles to ~24 (probably also 207) year cycles took place. This ‘event’ is synchronous with a change observed in Santa Barbara Basin varves, were interdecadal oscillatory modes are ‘turned on’ abruptly at that time. This might be taken as evidence for the hypothesis that varve formation in Lej da Silvaplauna responds to large scale atmospheric forcing. • It is suggested that during relatively colder intervals the oscillatory behavior of the forcing mechanisms is enhanced by larger catchment glaciers. During certain times decadal to centennial-scale oscillations are dampened or even rendered un- detectable because of the chaotic nature of the depositional system. First evidence for a non-linear response of varve formation to climatic forcing was presented. This is a reason for a word of caution when attempting to calculate transfer functions from a narrow ‘calibration’ time window chosen randomly from a larger dataset of climate ‘proxy’-data. Time series analysis and comparison with other climate proxy data clearly demonstrated that temporal changes in forcing factors

Chapter 7 115 Late Glacial in Lake Champfèr

CHAPTER 7

The Late- and Post-Glacial sediment record of Lej da Champfèr

7.1 Introduction

In view of possible future greenhouse warming Crowley and North (1991) write: “The rate of change of future global temperatures is comparable to or exceeds the greatest rates that occurred during the catastrophic deglaciations of the Pleistocene - the most extreme and abrupt, well-documented climate change recorded in the geologic record (cf. Jones et al., 1987)”. Therefore documenting the consequences and more importantly, understanding the processes that caused the “catastrophy” is perhaps the only chance to convince policy makers that it is dangerous to touch the dials of a very sensitive system. The aim of this chapter is to document how the last part of the deglaciation is recorded in the sediment sequence of Lej da Champfèr. In contrast to the moraine record, Lej da Champfèr sediments represent a continuous, high resolution record of the glacier history in the Upper Engadine. It is evaluated if any sediment properties can be identified that allow to trace the amount of catchment glaciation and the positions of the glaciers. Considering the results of chapter 4 it is assessed if also Late Glacial catchment runoff can be traced with an annual resolution. In this study the terms Late Glacial and Postglacial are used in a purely climatostratigraphic or better ‘glaciostratigraphic’ sense. This was chosen because sediment deposition in the Upper Engadine lakes is almost exclusively controlled by the degree of catchment glaciation (Leemann and Niessen, 1994b). The lake sedimentary records thus are interpreted solely in a process related way (presence or absence of glaciers in different parts of the catchment) without considering the timing of certain events. Terms like Holocene, Pleistocene or Younger Dryas and Preboreal, that are commonly used in chronostratigraphy, are avoided on purpose because dating of the sediments is not unambiguous yet. How the detected sedimentary changes could fit into the chronostratigraphic scheme is discussed in a separate section (7.3.6). Late Glacial in Lake Champfèr 116 Chapter 7

Questions that arose when Leemann and Niessen (1994b) presented the Postgla- cial sediment record of Lej da Silvaplauna, as a high resolution archive for the glacier history of the Upper Engadine were: (1) Is it possible to extend this record back into the Late Glacial ? (2) Can the same proxy indicators, Leemann and Niessen (1994b) used in the Postglacial sediments, i.e. variations in varve thickness and median grainsize, be applied to Late Glacial sediments in the same manner ?

PCH 94/3

N NW SE PSP 90/2 PSP 94/1 PCH 94/3

Postglacial Late Glacial

PSS 94/4

NW SE

PSS 94/4

10 ms ~ 7.3m

SW NE

PSP 90/2 PSP 94/1

Figure 7.1: 3.5 kHz seismic profiles recorded in lake Champfèr (northern basin), lake Silvaplana and lake Sils (southern basin). The Late-/Postglacial boundary is marked by an arrow in each profile. Core scetches of PSP 90/2 (Leemann, 1993), PCH 94/3 (cf. figure 7.2) and PSS 94/4 (cf appendix 7.1) are also plotted assuming a two-way travel time of 7.3 m/ms. Chapter 7 117 Late Glacial in Lake Champfèr

(1) In the initial phase of the project, it emerged from the seismic record (figure 7.1, Ohlendorf et al., 1995) that the northern basin of Lej da Champfèr and the southern basin of Lej da Segl (figure 2.6) are most promising, when Late Glacial sediments are to be cored. These basins are relatively protected and disconnected from direct glacial meltwater input today and most likely were during the entire Postglacial. Therefore, the Postglacial sediment infill in these basins is only around 5 m thick (figure 7.1), which represents a reduction by a factor of 2 with respect to Lej da Silvaplauna that receives glacial meltwaters today. This relatively thin Postglacial infill made it possible to recover 6.5 m and 5.2 m of Late Glacial sediments in Lej da Champfèr and Segl, respectively with the Kullenberg coring technique. This study will focus on the sediment record from Lej da Champfèr because it is the longest one and therefore probably offers the best time resolution and the oldest sediments. Another reason why high resolution sampling was restricted to this record is the fact that the sediments are continuously laminated throughout the entire 6.5 m of the Late Glacial section of the core (figure 7.2). This is not the case for the Lej da Segl sediments that show no lamination in the lowermost 1.6 m of the core (appendix 7.1). The Postglacial sediments of Lej da Champfèr are not the subject of discussion in this study since the Postglacial was investigated in detail in Lej da Silvaplauna (Leemann, 1993) and in Lej da S. Murezzan (Ariztegui, 1993). As mentioned above, the Postglacial in Lej da Champfèr basically represents a condensed version of the Lej da Silvaplauna Postglacial. The focus in this study is on the Late Glacial sediments, which have been analyzed in a high resolution mode. (2) Based on the analysis of Late Glacial varve thickness, median grainsize and mineral composition that is presented in this chapter the model developed in Leemann and Niessen (1994b) will be modified. It will be shown that for glacier size changes in a ‘cold climate mode’ like the Late Glacial additional parameters have to be considered that do not play a major role in a ‘warm climate mode’ for which the model was originally developed.

7.2 Results

7.2.1 Lithology of Lej da Champfèr sediments

The sediments of Lej da Champfèr recovered in the 1159 cm core PCH 94/3 (figure 7.2), that was taken at a waterdepth of 31 m in the northern basin of the lake (figure 2.6) have been divided into four lithostratigraphic units following Leemann Late Glacial in Lake Champfèr 118 Chapter 7

Lej da Champfèr (Northern basin)

Labels = Uncalibrated 14C-dates (years BP) Altitude : 1791m a.s.l. Opening date : 12.03.1996 Lines = Correlation between Water depth : 31m short-core and piston-core Core-length : 1159 cm FCH 94/1 PCH 94/3 121234567891011 9350

10 10

20 20

30 10370 30 10065

40 40 10330

50 50 7260

60 11315 60

70 70

80 80

11055 90 90 4728 17970

100 100

110 110

120 120

Figure 7.2: Photograph of gravity core FCH 94/1 and piston core PCH 94/3. The sediment/water interface is preserved in core FCH 94/1. Correlation of this core with the piston core is visualized by black lines. White numbers are uncalibrated 14C dates. Chapter 7 119 Late Glacial in Lake Champfèr and Niessen (1994b). Overall, the lithologic succession in core PCH 94/3 is very similar to core PSP 90/2 from Lej da Silvaplauna described in Leemann (1993). The main difference is the reduction of the Holocene sediment sequence from 10 m in Lej da Silvaplauna to only 5 m in Lej da Champfèr. A short sediment description and a representative photography of each unit and subunit are given in figure 7.3. Unit 1 (0-23 cm) comprises black and light grey, flocculent, laminated clayey silts with a high content of organic carbon (mean = 3.6%) and a high water content (mean = 72%). A peculiar feature of this unit is the occurrence of a 13 cm thick package of light grey, laminated sediment between 4 and 17 cm sediment depth. This package never shows any signs of internal deformation or erosive contact to the underlying, very soft sediments and has the same internal structure and thickness in different positions of the basin. It is sandwiched between black, organic rich, flocculent sediments with high water content above (0-4 cm) and below (17-23 cm). The black sediments reveal a rhythmic lamination when slightly oxidized and contain diatoms. This feature is observed in all cores taken from the northern basin of Lej da Champfèr, but in none of the other basins. Unit 2 (23-324 cm) is composed of laminated grey to brownish grey clayey silts and sands with irregular black mackles. The mean organic content of these sediments is 1.3% and with 62% the water content still is relatively high. A fine lamination becomes visible in some 1-5cm thick sections where black mackles are absent. According to the macroscopic inspection, this lamination is caused by differences in grainsize. Thicker coarser grained layers are overlain by thinner finer grained layers. Compared to the Lej da Silvaplauna, unit 2 varves (Leemann, 1993) the optical contrast between coarser and finer layers is much smaller. Therefore, some fine- grained sections that appear homogenous might in fact be laminated. The black mackles apparently are restricted to coarser layers, but they do not occur in every coarse layer. Due to their occurrence, preparation of the core is difficult especially when layers are bent, because they smear out and cover the surface with a black film that makes it impossible to see any lamination that might be present. Therefore, continuous, clearly rhythmic lamination, i.e. true varves, like in Lej da Silvaplauna unit 2 is not observed macroscopically in Lej da Champfèr unit 2 sediments. Unit 3 (324-505 cm) is characterized by laminated, black, grey, brown, greenish brown and white, clayey silts. Lamination is irregular but very fine. The organic carbon content of on average 3.9% is significantly higher than in unit 2, whereas the water content is only slightly lower (mean = 58%) but fluctuates a lot. At the top of unit 3 the transition to unit 2 is rather gradual. In contrast, at the bottom of unit 3 at 505 cm sediment depth, a very abrupt lithologic change to the underlying unit 4 is Late Glacial in Lake Champfèr 120 Chapter 7 *yr] 2 [mg/cm Accumulation rate 0.0 0.5 1.0 1.5 =1cm Foto of each unit [white bar = 1cm] Sediment description Black and light grey, flocculent, laminated, clayey silts Grey to brownish grey, laminated, clayey silts Black, grey, brown, greenish brown and white, finley laminated, clayey silts Greengrey, finley laminated, clayey silts Grey to brownish grey, coarsly laminated silts Grey to brownish grey, coarsly laminated silts with intercalated yellowbrown silt layers A green rhythmites B grey silts C yellow silts 1 2 3 4 4 4 5000 Dates 1200 2100 2740 4725 9350 10000 10065 10370 11315 11055 11000 10330 (12400) (12800) (13000) (13500) 7260 17970 C years BP 14 A B C 1 2 3 4 4 4 units Lithologic 0 1 2 3 4 5 6 7 8 9 10 11 [m] Sed. depth Figure 7.3: Core scetch of PCH 94/3. Definition of lithologic units follws Leemann (1993). The results of 14C datings are plotted in bold letters. Dates derived from the comparison of declination curves are shown in italics. Each lithologic unit is characterized by a short lithologic description and a representative photograph. A smoothed Chapter 7 121 Late Glacial in Lake Champfèr observed. This change is visible in all measured sediment properties (figure 7.4 and 7.5) and is immediately recognized in the sediments of all 4 Upper Engadine lakes (Ariztegui, 1993; Leemann, 1993, figure 7.2 and 7.3, appendix 7.1). It is described in more detail in section 7.3.1. Unit 4 consists of laminated grey, brownish grey, yellow-brown and green clayey silt and fine sand. In this unit the lowest organic carbon content of on average 0.16 ± 0.06% is observed. At the transition of unit 3 to 4, the water content drops abruptly to a mean of 26% in unit 4. Based on color and laminae thickness 3 different sediment types can be distinguished visually in unit 4. The first type (grey silts) is characterized by grey to brownish grey coarsely laminated silts. The second type (green rhythmites) is marked by green-grey, finely laminated, clayey silts. This type of lamination resembles most the rhythmic character of the Neoglacial varves from Lej da Silvaplauna (Leemann and Niessen, 1994b). The third type (yellow silts) is distinguished from the first one by the frequent occurrence of yellow-brown silt layers that are intercalated into the grey to brownish-grey coarsely-laminated silts. The yellow brown color often is restricted to the middle part of turbiditic layers. In unit 4 these three sediment types alternate several times from the top (505 cm) to the bottom (1159 cm) (figure 7.3). Therefore these sediments were further divided into 3 subunits (A-C) depending on the prevailing sediment type (figure 7.3). From the top to the bottom these are: Subunit 4A from 505 to 800 cm sediment depth consists mainly of green rhythmites with sections of greenish-grey silts at 505 - 580 cm and 610 - 665 cm. The lower green rhythmites (665-795 cm) of this unit show a homogenous bottom layer between 782 and 794 cm sediment depth. In the topmost part between 665 and 700 cm sediment depth lamination is coarser. Subunit 4B (grey silts) from 800 to 945 cm sediment depth mainly is composed of grey silts, but contains a section of yellow silts between 820 and 870 cm. Subunit 4C (yellow silts) from 945 to 1157 cm sediment depth is made out of yellow silts that are interrupted only by a very short section of green rhythmites at 989- 995 cm.

7.2.2 Physics, chemistry and mineralogy of the sediments 7.2.2.1 Lej da Champfèr core PCH 94/3 (0-1159 cm) The bulk mass accumulation rate (MAR) in Lej da Champfèr (figure 7.4a) varies between 0.1 and 0.5 g cm-2 yr-1 in units 1-3 with a minimum in the lower part of unit 3. At the unit 3/4 boundary that is encountered at 505 cm sediment depth the MAR Late Glacial in Lake Champfèr 122 Chapter 7 [SI] Magnetic susceptibility d 0 5 10 15 20 25 ] 3 c [g/cm Wet bulk density 0.8 1.2 1.6 2.0 [%] Water content b 10 30 50 70 90 *yr] 2 (MAR) [g/cm a 0.0 0.5 1.0 1.5 2.0 2.5 Mass accumulation rate B A C 1 2 3 4 4 4 units Lithologic 0 1 2 3 4 5 6 7 8 9 10 11 [m] Sed. depth Figure 7.4: Physical properties of core PCH 94/3. a) Mass accumulation rates (MAR), b) water content, c) wet bulk density and d) magnetic susceptibility. The supposedly warmest and the coldest interval on record are shaded in grey (see text for explanation). Heavy black line=smoothing with a 11 point (water content and susceptibility) and 31 point (MAR and Wet bulk density) running mean. Chapter 7 123 Late Glacial in Lake Champfèr increases by a factor of 10 from on average 0.1 g cm-2 yr-1 in unit 3 sediments to an average of 1.00 g cm-2 yr-1 in unit 4 sediments. This change is the most drastic and abrupt shift that is observed in the entire record. The water content (figure 7.4b) shows a gradual decrease from around 70% at the top of unit 1 to 60% at the bottom of unit 3. At the sediment surface and in the lower part of unit 3 a between 440 and 460 cm sediment depth (shaded in figure 7.4) local maxima are observed. At the unit 3/4 boundary the water content decreases abruptly to 30%. The wet bulk density (figure 7.4c) of the sediments is a mirror image of the water content. It also shows little variation in units 1-3 except for minima at the sediment surface and between 440 and 460 cm sediment depth. Wet bulk density increases abruptly at the unit 3/4 boundary from 1.3 g cm-3 in unit 3 to 1.8 g cm-3 in unit 4. The magnetic susceptibility (figure 7.4d) shows a fluctuation at the sediment surface in unit 1 and is very stable at a value of 10 SI in unit 2. In unit 3 it decreases to values between around 7.5 SI and reaches its minimum values of 3 SI between 440 and 460 cm sediment depth. In the lowermost part of unit 3 from 460 cm to the unit 3/4 boundary at 505 cm susceptibility values fluctuates considerably before it reaches stable values around 19 SI in unit 4. The organic carbon content (figure 7.5a) varies around 4% in unit 1 and remains relatively stable at 1.5% in unit 2. In unit 3 fluctuations are higher and the organic carbon content increases to maximum values of 7% between 440 and 460 cm sediment depth (shaded area in figure 7.5). At the unit 3/4 boundary a 16 fold decrease of the organic carbon content to the average unit 4 value of 0.16% is observed. When organic carbon accumulation rates are calculated (figure 7.5b) this decrease becomes less dramatic (3-fold). Unlike the organic carbon contents, the organic carbon accumulation rates do not change significantly from unit 2 to 3. Clearly the highest rates are observed in unit 1. Significant amounts of sulfur (figure 7.5c) were only detected in unit 1 and 3. The highest single value was measured at 395 cm sediment depth, but consistently high values are observed between 440 and 460 cm. In unit 4 sediments sulfur never was detected. The curve of carbonate carbon accumulation rates (figure 7.5d) shows the opposite trend than the sulfur curve. Virtually no carbonate is present in the unit 1-3 sediments. At the unit 3/4 boundary an abrupt increase to values around 6 mg cm-2 yr-1 is observed. Late Glacial in Lake Champfèr 124 Chapter 7 4 20 62 *yr] 2 12 81 d [mg/cm accumulation Carbonate carbon 04 2.0 3.0 [%] 1.0 Total sulfur 0.0 0 8 6 4 *yr] 2 81012 [mg/cm 62 41 21 Organic carbon accumulation 01 ab c [%] Organic carbon 01234567 A B C 1 2 3 4 units 4 4 Lithologic 0 1 2 3 4 5 6 7 8 9 10 11 [m] Sed. depth

Figure 7.5: Chemical properties of core PCH 94/3. a) Organic carbon content, b) Organic carbon accumulation rates, c) total sulfur content and d) carbonate carbon accumulation rates. The supposedly warmest and the coldest interval on record are shaded in grey (see text for explanation). Heavy black line=smoothing with a 11 point and 31 point running mean. Chapter 7 125 Late Glacial in Lake Champfèr

7.2.2.2 Magnetic susceptibility in other cores For interbasinal correlation magnetic susceptibility was measured in all cores that were taken from the 4 Upper Engadine lakes (figure 7.13 shows a selection). The susceptibility in core PCH 94/3 was described above. The downcore variation pattern is very similar in cores PSM 90/2 (Ariztegui, 1993) and PCH 94/4 (figure 7.13). The most prominent feature is the increase of susceptibility at the unit 3/4 lithologic boundary. In core PSS 94/4 this increase is much stronger. Susceptibility there increases 10-fold from 30 SI to 300 SI (figure 7.13). In this core susceptibility remains at this level until 850 cm sediment depth were a gradual decrease starts. Values around 60 SI are reached in 900 cm sediment depth. From there to the bottom of the core at 1080 cm susceptibility does not change significantly. The latter interval is marked by a different lithology (homogeneous, grey silty clay; cf. appendix 7.1) that does not occur in other cores. Except for the susceptibility record that is shown in figure 7.13 for the purpose of interbasinal comparison core PSS 94/4 is not discussed in this study. A correlation with the susceptibility record of core PSS 94/2 is not seen, because there is a large slump below 500 cm sediment depth.

7.2.2.3 Unit 4 sediments of PCH 94/3 (505-1159 cm) In the Late Glacial sediments of Lej da Champfèr changes of physical and chemical properties as well as changes in the mineral assemblage were determined with a time resolution of 2 to 10 years. Sampling was restricted to the non-turbiditic sediments. Bulk mass accumulation rates (MAR) were calculated from measured layer thicknesses assuming an annual nature of sediment deposition (see discussion under 7.3.5). Accumulation rates of organic and inorganic carbon are also based on this assumption. XRD peak intensities were taken as a semiquantitative measure for the content of the respective mineral in a given sediment sample (cf. 3.6)

Physical and chemical properties The median grainsize of unit 4 sediments (figure 7.6a) has proven to be a very robust parameter, because it can be determined with a standard deviation of usually less than 5%. For the fine grained Late Glacial sediments with an average median grainsize of 4.44 µm this implies that differences of >0.2 µm can be considered as significant variations. However, the variations that are detected in the median grainsize (figure 7.6a) also show up in the same manner when curves of the mean grainsize and the standard deviation (a measure of the degree of sorting) are plotted (appendix 7.2). The median of the Late Glacial sediments varies between 3.14 and 6.18 µm (figure 7.6a). Three sections of reduced median can be clearly detected (shaded Late Glacial in Lake Champfèr 126 Chapter 7 *yr] 2 MAR [g/cm 0.0 0.5 1.0 1.5 2.0 2.5 [SI] Susceptibility de 12 14 16 18 20 22 24 ] 3 density [g/cm Wet bulk 1.55 1.65 1.75 1.85 1.95 [%] Water content 20 25 30 35 40 [µm] median Grainsize abc 3.0 4.0 5.0 6.0 A B C 4 4 4 5 6 7 8 9 10 11 [m] Sed. depth Figure 7.6: Physical properties of the Late Glacial sediments in core PCH 94/3. a) Grainsize median, b) water content, c) wet bulk density, d) magnetic susceptibility. Sections of reduced median grainsize are highlited by vertical black bars. The supposedly coldest intervall is shaded in grey. Heavy black line=smoothing with a 5 (median), 11 (water content and susceptibility) and 31-point (wet bulk density and MAR) running mean. Chapter 7 127 Late Glacial in Lake Champfèr in figure 7.6). In the intervals 580-610 cm, 710-790 cm and 985-995 cm the median falls below the long term average of 4.44 µm to 3.70 ± 0.35 µm, 3.83 ± 0.36 µm and 3.63 ± 0.29 µm, respectively. Minimal values of <3.45 µm are always reached in these intervals. In all the other sections the median fluctuates around an average value of 4.7 ± 0.48 µm. Maxima of > 5.0 µm are observed at 550-570, 650-700, 800-940 and 1080-1090 cm sediment depth. Transitions between the sections of reduced median and sections with average median values or vice versa are always very abrupt. When compared to the lithologic description it becomes obvious that intervals of low median grainsize are confined to the green rhythmites of subunit A and C, whereas maximal median values occur in the grey silts of subunit A and B. The variations of the water content (figure 7.6b) are more or less a mirror image of the grainsize variations, i.e. higher water contents of up to 35% are observed for the sections of reduced median in the green rhythmites. The lowest water contents (19- 20%) occur at the lower and the upper boundary of the green rhythmites in subunit A. In the other sections the water content varies around an average of 25%. In the wet bulk density record (figure 7.6c) the green rhythmites show up as sections of reduced density. When fluctuations of the density are considered the lowest variability is noted in the lower grey and yellow silts of subunit B, whereas variability is highest in subunit A. The green rhythmites are also characterized by slightly higher susceptibilities (figure 7.6d). Values of up to 24 SI are observed, whereas in the other sections values are close to the Late Glacial average of 17 SI. This points to a slightly different mineral assemblage in the green rhythmites compared to the other two sediment types (cf. mineralogy). Bulk mass accumulation rates (MAR in figure 7.6e) vary between 0.14 and 4.57 g cm-2 yr-1 with an average of 1.00 g cm-2 yr-1. Minimal values occur in the green rhythmites of subunit A. Maxima are observed in the lower half of subunit C and in the yellow silts of subunit B. For the homogenous interval at the base of subunit A MAR values were calculated by interpolation. The organic carbon content (figure 7.7a) of the Late Glacial sediments is very low (mean= 0.16%). A linear increase is observed from the bottom of the core to middle of subunit B at 900 cm sediment depth. From there towards the unit 3/4 boundary the organic carbon content fluctuates between 0.1 and 0.25%. The only exception are maximum values of up to 0.4% that occur in the upper green rhythmites of subunit A. When organic carbon accumulation rates are calculated (figure 7.7b) a different picture emerges. Minimal accumulation rates are observed in subunit C and the lower green rhythmites of subunit A, whereas a broad maximum occurs in subunit B. Ac- Late Glacial in Lake Champfèr 128 Chapter 7 d *yr] 2 [g/cm accumulation Inorganic carbon 0 4 8 12 16 20 24 c [%] content Inorganic carbon 0.0 0.5 1.0 1.5 2.0 *yr] 2 [g/cm accumulation Organic carbon b 01234 [%] content a Organic carbon 0.05 0.15 0.25 0.35 A B C 4 4 4 5 6 7 8 9 10 11 [m] Sed. depth Figure 7.7: Chemical properties of the Late Glacial sediments in core PCH 94/3. a) Organic carbon content, b) organic carbon accumulation rates, c) carbonate carbon content, d) carbonate carbon accumulation rates. The supposedly coldest intervall is shaded in grey. Heavy black line=smoothing with a 5 point running mean. cumulation rates increase gradually in subunit A from 700 cm (top of the lower green rhythmites) towards a maximum at 580 cm (base of the uppermost grey silts). Differ- ent from most of the other measured parameters, transitions between maxima and minima of organic carbon accumulation are rather gradual. Chapter 7 129 Late Glacial in Lake Champfèr

The average inorganic carbon content (figure 7.7c) of the sediments is 0.7%. It was demonstrated (3.6 and figure 3.2) that inorganic carbon in fact represents detrital dolomite. Hence the average dolomite content is calculated to 5.4%. Maximum values are confined to subunit B and the lowermost homogenous part of subunit A. This picture does not change significantly when accumulation rates are calculated (figure 7.7d).

Mineralogy The mineralogy of the sediments is dominated by the allochthonous input of glacigenic rockflour. Autochthonous sediment components are either present in only minor amounts (see chapter 4) or completely absent in the Late Glacial sediments. Minerals that are present in the Late Glacial sediments with concentration of >5% were detectable by XRD analysis. In all subunits varying amounts of quartz, chlorite, mica, talc, plagioclase, orthoclase, serpentine, amphibole, pyroxene, dolomite and calcite were detected. Variations of the peak intensities of all detected minerals are shown in figures 7.9 to 7.11. A comparison of the intensities shows that quartz and phyllosilicates are the most abundant minerals. The phyllosilicates are approximately equal amounts of chlorites and micas. The latter are represented mostly by illite and muscovite, whereas biotite occurs 800 y = 0.8655x + 21.976 in only minor amounts. Listed by decreas- R2 = 0.7234 700 ing intensity plagioclase, dolomite and or-

(cps) 600 thoclase are also present in significant Å amounts in all samples. Minerals that do 500 not occur in all samples are amphibole, talc 400 and calcite. Serpentine was detected in the Quartz, 3.34 300 Late Glacial sediments in only 3 samples. 200 But in analogy to the sediments of Lej da 250 350 450 550 650 750 850 Mica (cps) Silvaplauna serpentine occurs more frequently in the early Postglacial sediments Figure 7.8: Scatter plot of mica 10.00Å versus quartz 3.34Å intensities. The good correlation in- (not shown). The quartz peak at 3.34 Å dicates that the quartz 3.34Å peak partly is over- shown in figure 7.9a can be viewed as a lain by a mica peak. mixed signal of the quartz and mica intensities. This becomes obvious from the good correlation of this peak with the mica 10.0 Å intensity (figure 7.8). Therefore the quartz peak at 4.26 Å is referred to as pure quartz. Several clear mineralogical signatures that are listed below were detected in the sediments. The clearest signature was detected for the green rhythmites. They are Late Glacial in Lake Champfèr 130 Chapter 7 [cps] Plagioclase 125 175 225 275 [cps] Orthoclase 25 50 75 bc d [cps] Quartz 4.24 Å 50 150 250 350 a [cps] Quartz 3.34 Å 250 450 650 850 A B C 4 4 4 5 6 7 8 9 10 11 [m] Sed. depth

Figure 7.9: Mineralogical properties of the Late Glacial sediments in core PCH 94/3. Intensities of the strongest characteristic peak for each mineral are plotted in counts per second (cps). a) Quartz (3.34Å), b) quartz (4.24Å), c) orthoclase (3.24Å), d) plagioclase (3.19Å). Note the low quartz intensities in the green rhythmites and low orthoclase intensities in the upper green rhythmites. The supposedly coldest intervall is shaded in grey. Heavy black line=smoothing with a 5 point running mean. marked by the lowest quartz (figure 7.9b) and mica (figure 7.10b) intensities the occurrence of amphibole (figure 7.11a) and talc (figure 7.11b) and the absence of calcite (figure 7.11c). Chlorite (figure 7.10a), plagioclase, orthoclase (figure 7.9c,d) and dolomite (figure 7.10d) intensities do not show a distinctive behaviour in the Chapter 7 131 Late Glacial in Lake Champfèr green rhythmites. The different mineralogy of the green rhythmites is best expressed as the chlorite/mica-ratio (figure 7.10c), because it pronounces the relative enrichment of chlorite. d [cps] Dolomite 0 100 200 300 c ratio Chlorite/Mica 0.5 0.8 1.0 1.3 1.5 b [cps] Mica 250 350 450 550 650 750 a [cps] Chlorite 350 450 550 A B C 4 4 4 5 6 7 8 9 10 11 [m] Sed. depth Figure 7.10: Mineralogical properties of the Late Glacial sediments in core PCH 94/3. a) Chlorite (7.00Å) intensities, b) mica (10.00Å) intensities, c) chlorite/mica ratio, d) dolomite (2.89Å) intensities. Note the high chlorite/mica ratios in the green rhythmites and the high dolomite intensities in unit 4B. The supposedly coldest intervall is shaded in grey. Heavy black line=smoothing with a 5 point running mean. Late Glacial in Lake Champfèr 132 Chapter 7 0 06 [cps] Calcite 04 02 0 06 Talc [cps] 04 02 0 06 ab c [cps] 04 Amphibol 02 A B C 4 4 4 5 6 7 8 9 10 11 [m] Sed. depth Figure 7.11: Mineralogical properties of the Late Glacial sediments in core PCH 94/3. a) Amphibole (8.45Å) intensities, b) talc (9.31Å) intensities, c) calcite (3.03Å) intensities. Note the occurrence of amphibole and talc and the absence of calcite in the green rhythmites. The supposedly coldest intervall is shaded in grey. Heavy black line=smoothing with a 5 point running mean.

As described for the median grainsize curve, also the change in mineralogy are partly abrupt. The abruptness may be exemplified by the clear shift between the upper green rhythmites and the grey silts of subunit A at 580 cm sediment depth. The upper green rhythmites, in addition to the characteristic mineralogy described above, show the lowest orthoclase intensities (figure 7.9c). Immediately above the Chapter 7 133 Late Glacial in Lake Champfèr boundary at the base of the topmost grey silts high orthoclase, plagioclase and quartz intensities and low chlorite and dolomite intensities are observed. There, calcite occurs simultaneously with low dolomite intensities. Another shift in sediment mineralogy that is more gradual and not as dramatic as the one described above occurs at the subunit B/C boundary. There the chlorite intensities (figure 7.10a) decrease from 520 cps in the yellow silts (C) to minimal intensities of 350 cps at 910 cm sediment depth in the overlying grey silts (B). From this minimum towards the top of subunit B chlorite concentrations increase again gradually. At the same boundary also the dolomite intensities (figure 7.10d) increase sharply and show a peak at 930 cm sediment depth. Higher dolomite intensities are in general characteristic of subunit B. Plagioclase and orthoclase (figure 7.9c,d) show peak intensities a little later than dolomite, at 910 cm sediment depth, exactly were minimal chlorite intensities were detected. Plagioclase intensities decrease dramatically afterwards to reach their minimum values in the yellow silts of subunit B. In summary, considering all measured parameters, the green rhythmites stand out as a sediment type with distinctly different physical, and mineralogical properties. In particular they are characterized by low median grainsize, MAR, quartz and mica values and high water content and chlorite/mica-ratio as well as by the occurrence of amphibole and talc but the absence of calcite.

7.2.3 Dating of Lej da Champfèr sediments

For the Holocene sediments a chronology is well established for adjacent Lej da Silvaplauna (figure 7.12c) which is based on AMS 14C-dating, varve counts and declination measurements (Leemann and Niessen, 1994b). This chronology was transferred to Lej da Champfèr Holocene sediments by correlating declination curves and characteristic lithologic boundaries (figure 7.12c). It is constrained by one AMS 14C-date of 4725 BP (figure 7.3 and 7.12a). Only between 340 and 440 cm sediment depth the declination swings are not very pronounced and the intensity of the palaeomagnetic vector is not very high (cf. appendix 7.3). Therefore, a correlation of declination swings is ambiguous in this interval. But for the other Holocene parts of the core the correlation with Lej da Silvaplauna declination swings b, c and d is very obvious and the lithologic succession is similar to that in Lej da Silvaplauna (figure 7.12a-c). Therefore no further AMS dating was done on the Holocene sediments. Instead, all dating attempts were restricted to the pre-Holocene sediments below 505 cm sediment depth (see below). The Lej da Champfèr declination curve in figure 7.12c, shows regular fluctuations that are typical for secular variation patterns in the Late Glacial sediments, too. They Late Glacial in Lake Champfèr 134 Chapter 7 d b f PSP 90/1 [ ] e c Leemann, 1993 -30 -10 10 30 Declination b d ? PCH 94/3 c Θ [ ] e I-j? I-l? c 60 120 180 240 300 360 I-k? Declination I? I-m? This study 16000 C-dates 14 C years BP 14 errors σ in identified Macrofossils and extrapolation downcore with the same trend. Linear interpolation between 2 AMS b Calibrated C years BP) C years BP with 2 14 14 2000 4000 6000 8000 10000 12000 14000 Calibrated Palaeomagnetic declination with estimated errors (calibrated Lithologic boundary at 9350 BP (AMS-dating, this study) Lithologic boundary at 10000 BP (Declination data, Pollen data from e.g. Wick & Tinner, 1997) 0 5000 BP 1200 2100 2740 4725 9350 10000 10065 10370 11315 11055 11000 10330 (12400) 7260 17970 (12800) (13000) (13500) C years Dates a 14 A B C 1 2 3 4 4 4 logic units Litho- 1 2 3 4 5 6 7 8 9 0 10 11 [m] Sed. depth Figure 7.12: Different possible chronologies for the Late Glacial sediments of core PCH 94/3. a) core sketch and lithologic units. 14C-dates shown in bold, declination dates in italics. b) Calibrated (Stuiver and Reimer, 1993) 14C-ages plotted versus sediment depth. Heavy and thin black lines for the Late Glacial sediments are derived from layer counting, assuming annual deposition. Heavy grey line is derived from extrapolation of the 14C-dates at 9350 and 10065 BP (see text for explanation). c) Declination swings of PCH 94/3 compared to declination swings in core PSP 90/2 (Leemann, 1993). Swings are labeled after Leeman (1993) and Wessels (1995). Chapter 7 135 Late Glacial in Lake Champfèr are difficult to interpret because apparently rotation of the core barrel during pen- etration of the sediments lead to a systematic drift of declination swings. Values decrease from 300° at the top of the core to 60° at the bottom. Also a rotation of the sediment with respect to the core-liner during measurement can not be excluded. Therefore, only the two clearest declination swings are labelled in figure 7.12c. Labelling of the Late Glacial sediments follows Wessels (1995), whereas the Holocene labels are adapted from Leemann and Niessen (1994b). The swing labelled with Q in figure 7.12c is relatively well dated in lakes Zürich (Lister, 1985), Lugano (Niessen and Kelts, 1989) and Constance (Wessels, 1995) to 10000 BP. The next relatively large swing (I) was dated to 11000 BP by Wessels (1995). For the Late Glacial sediments of Lej da Champfèr in total 16 samples were taken for AMS 14C-dating. The mean organic carbon content of these sediments is very low (0.16 ± 0.06%). Hence, from these samples, 8 contained too little organic matter for dating and 8 were dateable. The dates obtained are given in table 7.1 and figure 7.12a. Among these 8 dates the date from the base of the core (sample 10 in table 7.1), due to the absence of macrofossils, was done on bulk organic matter. The relatively high d13C-value of this sample indicates a dominant contribution of aquatic plants. 14 Aquatic plants are incorporating CO2 that is depleted in C because it originates from mesozoic carbonatic rocks in the catchment or from the meltwater of old glacier ice (Heitz et al., 1982b). The resulting AMS date of 17970 BP (table 7.1) thus is very likely to be too old and is disregarded for this reason.

Table 7.1: AMS 14C-dates determined in the Late and Postglacial sediments of core PCH 94/3 from Lej da Champfèr.

δ 13C value Age in Section Sediment (‰) AMS 14C calibrated Sample Core/Section depth depth [Error: ± age BP Error years BP No. Lake No. (cm) (cm) Lab No. Material dated 1.0-4.0 ‰] (yrs) (yrs) (yrs) 1b Champfèr PCH 94/1-4 36.5-37.5 390.0 ETH-15099 Picea needle, cat scale -25,2 4725 ± 80 5452 5611 - 5294 Picea needle, cat scale, small 2 Champfèr PCH 94/1-5 43.5-44.0 506.0 ETH-15100 twig, woody fragment -26,2 9350 ± 90 10299 10554 - 10044 small twigs, needle, bork 3 Champfèr PCH 94/1-7 40.0-41.0 639.0 ETH-15101 fragment -26,4 10065 ± 90 11565 12129 - 11000 unidentified, black plant 18 Champfèr PCH 94/3-7 31.0-32.0 737.0 ETH-17519 fragments -21,9 10370 ± 115 12257 12560 - 11715 unidentified, black plant 17 Champfèr PCH 94/3-7 58.0-59.5 765.0 ETH-17518 fragments -25,0 11315 ± 185 13222 13667 - 12851 unidentified, black plant 16 Champfèr PCH 94/3-8 84.0-87.0 891.0 ETH-17517 fragments -21,5 11055 ± 110 12968 13211 - 12736 unidentified, black plant 15 Champfèr PCH 94/3-9 43.5-45.5 950.0 ETH-17516 fragments -21,9 10330 ± 130 12199 12547 - 11446 10 Champfèr PCH 94/3-11 88.0-94.0 1149.5 ETH-16516 bulk -15,4 17970 ± 190 21445 22023 - 20827 Fragments of a terrestrial 19 Champfèr PCH 94/3-11 53.0-55.0 1114.0 UtC 6761 beetle -8,0 7260 ± 300 8046 8569 - 7523 Another date close to the base of the core (sample 19 in table 7.1) was done on the remains of a terrestrial beetle and a few needle fragments. Since this sample was extremely small (it yielded only 20µg of carbon) the danger of contamination during the sampling procedure is very high. The very high d13C-value of -8 ‰ also indicates a severe contamination. The age obtained thus is much too young and is also rejected. Late Glacial in Lake Champfèr 136 Chapter 7

All the other dates were determined on terrigenous plant remains as indicated by the d13C-values. Only in two cases the plant remains could actually be identified (sample 2 and 3 in table 7.1). These two dates are considered the most reliable ones. Since the Late Glacial sediments are continuously laminated (cf. 7.2.1 and figure 7.2) the number of layer couplets that are present in the 650cm of Late Glacial sediments was counted. In total 2600 layer couplets were counted between 505 and 1159cm sediment depth. Assuming the couplets are varves this would equal to 2600 years. The result of the layer counting is presented in figure 7.12b. This floating varve chronology is tentatively anchored at different absolute ages (see 7.3.4 and 7.3.5 for discussion).

7.3 Discussion

7.3.1 Comparison of Late Glacial and Postglacial sediments

This section serves to demonstrate the tremendous difference between the Post- glacial, unit 1-3 and the Late Glacial, unit 4 sediments in the Lej da Champfèr record (figures 7.4 and 7.5). In Lej da Silvaplauna Leemann and Niessen (1994b) ascribe the 9-times higher MAR of the late Postglacial as compared to the early Postglacial, to increased glacial erosion with the advent of Neoglaciation in the catchment around 3300 BP. They postulate no, or only negligible glacial activity in Lej da Silvaplauna catchment during the early and middle Postglacial. In the Lej da Champfèr core (PCH 94/3) the Neoglacial (unit 2) MAR are only 3.5 times higher than early and middle Postglacial (unit 3) MAR (figure 7.4a). The lowest MAR (shaded area in figure 7.4a) in Lej da Champfèr, like in Lej da Silvaplauna, are observed in the lower part of unit 3 (early Postglacial). This is also the only sediment section were significant amounts of sulfur could be detected (figure 7.5c). As mentioned in chapter 5, this is supposedly due to the fixation of sulfur in authigenic iron sulfides in sub- or anoxic bottom waters at that time. This would point to sluggish water circulation and/or higher productivity during this interval. Presumably this section represents the warmest time period on record. In the same interval (440 and 460 cm) a wet bulk density (figure 7.4c) minimum occurs that coincides with maxima in the water content (figure 7.4b) and the organic carbon accumulation rates (figure 7.5b). The good inverse correlation of wet bulk density with the water content (r2=0.92) and the organic carbon content (r2=0.59) indicates that density variations are to a large extent controlled by changes in these Chapter 7 137 Late Glacial in Lake Champfèr two parameters (e.g. Menounos, 1997). Especially the good correlation with the water content shows, that wet bulk density variations actually are a measure for the porosity of the sediments. Going back into the Late Glacial, the switch over between a heavily glaciated (Late Glacial) and glacier free (Postglacial) catchment should be even more dramatic than the onset of Neoglaciation as indicated by the abrupt transition from unit 3 organic rich laminites to unit 4 ‘rhythmites’ in the lowermost 50 cm of the Lej da Silvaplauna core (Leemann and Niessen, 1994b). In Lej da Champfèr this unit 3/4 boundary is encountered at 505 cm sediment depth (figure 7.3). It is the most drastic and abrupt lithologic change (10-fold MAR increase) of the entire record (figure 7.2 and 7.3). Downcore, at this boundary the water content, the organic carbon and the total sulfur content decrease whereas MAR, wet bulk density, magnetic susceptibility and carbonate carbon accumulation rates increase (figure 7.4 and 7.5). The most marked difference between the unit 4 and unit 1 to 3 sediments is the almost complete absence of carbonate in the Postglacial sediments (figure 7.5d). Since it was demonstrated that virtually all carbonate is detrital dolomite (3.6), this indicates that the dolomite source for Lej da Champfèr was ‘switched off’ at the end of the Late Glacial and not ‘switched on’ again during the Postglacial. This seems somewhat surprising, considering the fact that detrital input increased again (although not as strong as in Lej da Silvaplauna) with the onset of Neoglaciation (figure 7.4a). Therefore, detrital carbonate in the Late Glacial sediments presumably was carried to Lej da Champfèr mainly by meltwaters from the glaciers at the western slope of Piz Corvatsch because only there carbonates outcrop in significant amounts (figure 2.3 and 7.14). Carbonate free Postglacial sediments indicate that Neoglaciation of this part of the catchment was small compared to Late Glacial advances. In summary, following Leemann and Niessen (1994b) the very drastic sedimentological change at the unit 3/4 boundary is viewed as the complete cessation of glacial meltwater input into Lej da Champfèr. As a consequence of this, accumulation rates decrease rapidly by a factor of 10 (figure 7.4a). This again underlines the enor- mous importance of glacial erosion for sediment accumulation. It is referred to subsequently as the Late Glacial/Postglacial boundary.

7.3.2 Is the unit 3/4 boundary an isochrone ?

The boundary between unit 3 and 4 is clearly detectable in all 4 Upper Engadine lakes (figure 7.13). Subsequently it is discussed whether this boundary represents an isochrone. Late Glacial in Lake Champfèr 138 Chapter 7 120 8880 ± Lej da 01030 50 PSM 90/2 S. Murezzan 0 1 2 3 4 5 6 7 8 9 10 11 12 Ariztegui (1993) 0 5 10 15 20 25 PCH 94/3 N 0 1 2 3 4 5 6 7 8 9 10 11 This study Lej da 90 ± 9350 Champfèr 0 10203040 PCH 94/4 0 1 2 3 4 5 6 7 8 9 This study Lej da Silvaplauna 0 100 200 PSP 90/2 0 1 2 3 4 5 6 7 8 9 10 11 Leemann (1993) 90 ± 9415 Post Late Glacial Glacial 0 100 200 PSS 94/2 ? ? 0 1 2 3 4 5 6 7 8 This study Segl Lej da SN S 0 100 200 300 PSS 94/4 0 1 2 3 4 5 6 7 8 9 10 This study

Figure 7.13: Core sketches and magnetic susceptibility profiles of 6 piston cores from all basins of the Upper Engadine lakes. 5 cores show a lithologic succession similar to that of core PSP 90/2 described in Leemann (1993). The unit 3/4 boundary (dashed grey line) is clearly recognized in these cores. 3 AMS 14C-dates on terrestrial macro remains are also shown. Note the younger age determined in Lej da S. Murezzan (see text for discussion). Chapter 7 139 Late Glacial in Lake Champfèr

The 4 Upper Engadine lakes can be divided into 6 different basins (figure 2.2). From these, 2 basins, i.e. the northern basin of Lej da Segl and Lej da Silvaplauna, receive direct glacial meltwater input today (figure 2.2 and 2.4). Meanwhile long piston cores are available from all of these basins (figure 2.6 and 7.13, Ariztegui, 1993; Leemann, 1993). The lithologic succession comprising the 4 lithologic units described above is very similar in all 4 lakes. It is also reflected in the physical properties of the sediments. Magnetic susceptibility is plotted as an example in figure 7.13. Especially the very marked unit 3/4 boundary which was characterized in the previous section for the northern basin of Lej da Champfèr, is easily recognized. This boundary was cored in 5 of the basins (figure 7.13). Only in the northern basin of Lej da Segl it was not possible to penetrate this boundary. The Postglacial sedimentary infill of this basin is much thicker than in all the other basins for various reasons. First of all, the glacial river (Aua da Fedoz) feeds directly into this basin. There is no morphological barrier between the large Isola delta built by the glacial river and this basin (figure 2.6). In addition, the adjacent slopes of this basin are very steep and often give rise to avalanches and debris flows. Consequently the sediments of this basin basically consist of a sequence of homogenites and turbidites with intercalated short sections of laminated silts. In all the other basins the unit 3/4 boundary was cored. Figure 7.13 shows one core from each basin and its respective magnetic susceptibility variations that serve as an additional correlation tool. The unit 3/4 boundary is highlighted by a heavy grey line. It is noted that the susceptibility increases abruptly at this boundary in all basins but at different magnitudes. The Postglacial sediments all show similar susceptibilities between 10 and 20 SI for all the basins. This is not the case for the Late Glacial sediments. Lej da Champfèr and S. Murezzan show 2 times higher susceptibilities in the Late Glacial sediments, whereas in Lej da Segl and Silvaplauna Late Glacial susceptibilities are more than 10 times higher. The moderate increase for the former can be explained by the higher density of the Late Glacial sediments (figure 7.4) since susceptibility is volume dependent, which was not accounted for in the raw data shown in figure 7.13. Hence, the input of magnetic minerals probably did not change with deglaciation of the catchment. In contrast, the significantly stronger increase of susceptibility in the other two lakes can not be explained by a change in sediment density alone. This indicates that the species, the grain size and/or the amount of magnetic minerals that were introduced into the lake was different during the Late Glacial. It is therefore proposed that extended Late Glacial glaciers eroded lithologies that are enriched in minerals with high specific susceptibilities. These most likely are the Platta and Forno-Lizun ophiolites that outcrop at the western slopes of Piz Corvatsch, at Grevasalvas and in Late Glacial in Lake Champfèr 140 Chapter 7

Val Forno (figure 2.3 and 7.14). Magnetite, titanite, ilmenite and hematite e.g. have been described by Finger et al. (1982) in greenschists of the Platta ophiolites. Due to their higher density and/or coarser grainsize, these minerals were deposited only in the basins that received direct meltwater input from this catchment. The downstream basins of Lej da Champfèr and S. Murezzan always only received suspension where magnetic minerals were filtered out by the upstream basins. The consistently low Postglacial susceptibilities therefore indicate that the respective rocks were never eroded during the Neoglacial glacier advances. From the interbasinal comparison of sediment thicknesses (figure 7.13) it is clearly seen that the Postglacial sedimentary infill decreases with increasing distance from the direct glacial meltwater input, i.e. from the northern basin of Lej da Segl downstream to the northern basin of Lej da Champfèr. The southern basin of Lej da Segl also is presently disconnected from direct meltwater input. Lej da S. Murezzan is, in addition, influenced by glacial meltwaters from Vadret Güglia fed into the En by the Ova da Suvretta (figure 7.14) and thus shows a thicker Postglacial infill than the very protected northern basin of Lej da Champfèr. In three of the basins the unit 3/4 boundary was dated by AMS 14C on terrigenous macro remains (Ariztegui, 1993; Leemann, 1993) and this study). All three dates were determined on samples that were taken directly at the boundary and not below or above it. The dates are shown in figure 7.13. It emerges that the 14C ages of the boundary are becoming successively younger from Lej da Silvaplauna to Lej da S. Murezzan. Whereas the Lej da Champfèr and Lej da Silvaplauna dates overlap within the limits of error and therefore can be considered as isochronous, the Lej da S. Murezzan date clearly is ~500 years younger. If this is not an artefact caused by the very low accumulation rates at the base of unit 3, the question concerning the diachronism arises. It was assumed above that this drastic lithologic change is caused by the cessation of glacial meltwater input. Accepting this assumption, this would indicate that Lej da S. Murezzan still received glacial meltwater until 8880 BP, while the Lej da Silvaplauna and Champfèr catchment already was deglaciated around 9400 BP. There are three possible scenarios that could explain this lag time of more than 500 years. 1. The small glaciers at the northeast facing slope of Piz Güglia (Vadret Güglia) and the north facing slope of Piz Rosatsch (Vadret da Rosatsch) persisted for a longer time than the glaciers in Val Fex and Val Fedoz (figure 2.1 and 7.14). The former two glaciers only feed into Lej da S. Murezzan and do not affect the other 3 lakes to the southwest. 2. A hydrological connection of Lej da S. Murezzan to the Flaz system that drains the large glaciers of the Bernina group, e.g. Vadret da Morteratsch and Vadret da Roseg (figure 2.4) still existed until 8880 BP. Chapter 7 141 Late Glacial in Lake Champfèr

3. A body of dead ice persisted much longer in the Lej da S. Murezzan basin than in the other lake basins as has been proposed by Kleiber (1974). She even suggested that the dead ice might have persisted until 7500 BP. Since this interpretation is based on bulk 14C dating of gyttja from a core taken at Mauntschas, a pond close to the eastern shore of Lej da S. Murezzan (figure 2.4), it has to be questioned. With the available data a decision in favour of one of these three possibilities can not be made. But considering the differences in mineralogy of the catchment rocks (figure 2.3 and 7.14) it should be possible to identify rock flour derived from the Bernina rocks in the Late Glacial sediments of Lej da S. Murezzan by means of XRD- analysis. A long sediment core from Lej Staz (figure 2.1 and 2.4), a small shallow lake draining into Lej da S. Murezzan from the east, could also help to further elucidate the latest phase of deglaciation and a possible connection to the Flaz system. Re- dating of the pollen profiles (figure 2.4) that were measured in the 1970s and 80s from several small ponds in the vicinity of the Upper Engadine lakes (e.g. Lej Marsch, Mauntschas, Maloja) with modern techniques would be an important task too. Fur- thermore 14C-dating of the unit 3/4 boundary in Lej da Segl (southern basin) would yield valuable information about the timing of deglaciation in the southern part of the catchment especially in Val Forno. In conclusion, the susceptibility records suggest that (1) the glacier erosion of magnetic mineral-bearing rocks in the Segl and Silvaplauna catchment stopped at the LGl/PGl boundary. (2) The Lej da Champfèr and S. Murezzan susceptibility records are not influenced by this change. The comparison of 14C dates for the unit 3/4 transition in 3 of the 4 Upper Engadine lakes suggests that the cessation of glacial meltwater input might not have been synchronous in all lakes. Consequently this boundary should not be used as a chronostratigraphic reference horizon. How- ever, to verify this hypothesis it is necessary to conduct high resolution AMS-14C dating across the unit 3/4 boundary in all 4 lakes. Moreover, the timing for the development of vegetation with respect to the disappearance of the catchment glaciers is unclear yet. For Lej da Champfèr these questions are addressed in an ongoing study (E. Gobet, pers. comm.).

7.3.3 Special characteristics of the green rhythmites

It was mentioned above that wet bulk density can be regarded as a measure for the porosity of the sediments. From this it was expected that in the Late Glacial sediments the intervals of coarser grain size (figure 7.6a) would show lower wet bulk densities (figure 7.6c) as a result of higher porosity and thus higher water content. It was somewhat surprising that the opposite is the case, i.e. the fine grained green Late Glacial in Lake Champfèr 142 Chapter 7 rhythmites show lower wet bulk densities and higher water contents (figure 7.6a-c). Menounos (1997) observes the same phenomenon in an high alpine lake (Sky Pond) in the Rocky mountains. A possible explanation for this feature could be a relatively higher content of phyllosilicates that might increase the porosity by forming ‘card- house structures’. According to F. Niessen (pers. comm.) low wet bulk densities of Late Glacial marine sediments from the Arctic Ocean are the result of a flocculation of fine-grained particles in the water column and their sedimentation as aggregates which then create the higher porosity, i.e. lower density. This might be another explanation for the observed density variation but it remains to be tested, whether this model also applies to freshwater environments. However, a smaller median grainsize does not necessarily imply a higher propor- tion of phyllosilicates, since it has been shown that the glacial grinding process can produce quartz grains of less than 2 µm (Small, 1987b). The XRD analysis revealed that quartz (figure 7.9b) as well as mica intensities (figure 7.10b) are lower in the green rhythmites whereas chlorite intensities (figure 7.10a) are not, although they fluctuate considerably. Over all, this results in a relative enrichment of chlorite in the green rhythmites (figure 7.10c). The only other minerals that are enriched are talc and amphiboles (figure 7.11a,b). It is therefore possible that the relative enrichment of platy (chlorite, talc) and stalked (amphibole) minerals is responsible for the observed lower median grainsize and thus wet bulk density of the green rhythmites. In addition the relative enrichment of chlorite (figure 7.10c) together with the low quartz intensities (figure 7.9b) and the exclusive occurrence of talc (figure 7.11b) might be responsible for the green color of these sediments. Furthermore, the enrichment of amphiboles (figure 7.11a) could be responsible for the magnetic properties of the green rhythmites, because amphiboles are paramagnetic minerals. Their specific susceptibility of 10-100 * 10-8 m3 kg-1 hence is significantly higher than that of all the other minerals that were detected in the sediments by XRD analysis. Therefore it is likely that their occurrence causes the higher susceptibility of the green rhythmites (figure 7.6d).

7.3.4 Median grainsize and MAR: Glacier size proxies in the Late Glacial sediments of Lej da Champfèr ?

Following the model of Leemann and Niessen (1994b) in a system were glacier- derived and non-glacial, terrigenous detritus are mixed, a reduction of the sediment median grainsize should be indicative of a more extensive catchment glaciation, that produces large amounts of fine grained rock flour. Leemann and Niessen (1994b) suggested this glacier size proxy to distinguish between extreme situations (end Chapter 7 143 Late Glacial in Lake Champfèr members) in the Holocene of Lej da Silvaplauna. There the early Holocene sediments deposited with no catchment glacier cover have a median grainsize of 16 µm on average. This represents pure background sedimentation, i.e. terrigenous detritus that is washed into the lake by rainfall. The difference to the late Holocene sediments (median grainsize = 6.5µm), where the terrigenous detritus is strongly ‘diluted’ by glacier-derived debris, is quite large. Compared to this, grainsize fluctuations within the Neoglacial period are relatively small (± 1.25 µm, Leemann, 1993). Therefore, the authors suggest to use the sediment accumulation rates (MAR) as proxy for glacier size variations during the Neoglacial on decadal to centennial scales (cf. chapter 4). High MAR then indicate more input of glacigenic rockflour as a result of an increased erosion capacity during glacier high stands. In contrast, low MAR characterize times with reduced catchment glaciation. Ariztegui et al. (1997) demonstrated that even small variations in the median grain size during times of persistent glacier cover can serve as tracers for glacier size variations or changes in non-glacial sediment input. Ariztegui et al. (1997) were able to show that the Younger Dryas cold interval can be recognized in the sediments of lake Mascardi (Argentina) as a relatively small reduction of the median grainsize from ~5.5 µm in the Allerød to ~4.5 µm during the YD glacier highstand. This interpretation is supported by pollen and geochemical evidence. Hence, in analogy to this, in the Late Glacial of Lej da Champfèr the minimal median grainsize determined for the green rhythmites is viewed as evidence for their deposition during a glacier highstand. Given the hypothesis of Leemann and Niessen (1994b) outlined above, times of glacier highstands should also result in higher MAR. From figure 7.6e it becomes obvious that the opposite is the case for the Lej da Champfèr sediments. There, the grainsize minima of the green rhythmites are accompanied by minima of MAR. Due to this contradiction it has to be concluded that the model of Leemann and Niessen (1994b) cannot be applied in the same way to the Late Glacial sediments when the catchment glaciation was significantly higher. The amount of non-glacial sediment components that could modulate the grainsize variations was strongly reduced compared to Neoglacial times. Therefore, in contrast to Postglacial times, grainsize variations in the Late Glacial sediments most likely are predominantly caused by glacier-derived debris. Glaciologic processes thus are much more important. This is a fundamental difference to the record of Leemann and Niessen (1994b). Their model therefore has to be modified. One possible explanation for the observed sediment properties might be a different mode of glacier movement during the Late Glacial. The simultaneous occurrence of a reduced median and low MAR could be the result of a lower mass balance Late Glacial in Lake Champfèr 144 Chapter 7 gradient and lower basal shear stresses of the catchment glaciers. This would point to a cold and arid climate (Maisch and Haeberli, 1982). Perhaps glaciers were even partly frozen to the ground which would reduce their sediment production dramatically. But even if the glaciers were not cold based, for an advancing glacier with a positive mass balance, a reduced production of meltwater has to be assumed, since otherwise it would not advance. Reduced meltwater production (=reduced runoff), in turn, results in reduced transport capacity of the glacial outwash stream. In the receiving lake this will lead to lower accumulation rates of finer grained sediments as observed for the green rhythmites. Accordingly, a retreating glacier with a negative mass balance releases more meltwater and suspension resulting in a higher transport capacity of the glacial outwash stream and perhaps reworking of subglacial sediments. Consequently, sediments deposited in the receiving lake will show higher accumulation rates and coarser grainsize. For equilibrium conditions i.e. a glacier remains stable in an advanced or retreated position (mass balance near 0) two extreme modes can be distinguished. (1) the glacier has a high mass balance gradient. In this mode it is very active, i.e. basal shear stress is very high and rocks at the glacier bed are eroded intensively. The suspension load and transport capacity of the glacial outwash stream are high. (2) the glacier has a low mass balance gradient. Then, the glacier is rather inactive and basal shear stress is low. Erosion of rocks at the glacier bed is small or is not taking place at all, if the glacier is frozen to the ground. Only very little suspension is produced and transport capacity of the glacial outwash stream is low. The first mode is characteristic for a maritime climate with high precipitation, e.g. glaciers at the western coast of Scandinavia or North America, whereas the second mode usually occurs in a cold, continental climate with low precipitation, e.g. glaciers in the Canadian Arctic or central Asian mountains, (Haeberli, 1994). Consider- ing these two end members the basal shear stress of a glacier, is a measure for the degree of continentality of local climate. Although the basal shear stresses might be the main control mechanism for sediment deposition in Lej da Champfèr it has to considered that there are numerous other factors that can influence the grainsize distribution in the glacier milk and the amount of fine-grained rockflour available for deposition in the proglacial lake (Haeberli, 1994; Haeberli, 1995; Small, 1987a; Souchez and Lorrain, 1987), e.g.: • Does the glacier movement take place in a sediment bed (eroding it’s own till) or over bare rock ? • The lithology underneath the glacier. • The presence and concentration of debris in the basal ice layer. Chapter 7 145 Late Glacial in Lake Champfèr

• Do most of the particles originate from subglacial or from englacial sources ? It was shown that only the subglacial grinding process can produce large amounts of fine grained rockflour (Small, 1987a). • Does the glacier advance over pre-existing stagnant ice ? If this is the case subglacial erosion is at a minimum and takes place only at the very glacier front (Habbe, 1996). • It has to be considered that Lej da Champfèr mainly received ‘filtered’ meltwater. Thus the availability of fine grained rockflour rather than the transport capacity of the glacial outwash stream might be the critical factor controlling sediment grainsize and MAR. It is not known how important the influence of these interfering processes is, with respect to the concept of basal shear stresses presented above. Nevertheless, the ‘basal shear stress model’ is considered as a valid simplification. In particular, the reduction of the sediment grainsize in combination with the observed mineralogy can still be taken as evidence for maximum glacier extension, because both are directly related to subglacial erosion. For the Engadine Maisch and Haeberli (1982) reconstructed increasing basal shear stresses of glaciers during the Late Glacial from palaeo-glacier geometries. From this observation the authors inferred that Late Glacial climate warming was accompanied by an increase in precipitation. Given an intensive glaciation in the catchment of the Upper Engadine lakes during the Late Glacial it is assumed that the basal shear stress is an important parameter that controls the amount of suspension released. High MAR and median grainsize hence can be the result of both, increased basal friction or retreat of the glacier with reworking of subglacial deposits. In the following the interpretation of the sedimentary record of Lej da Champfèr is based on this hypothesis. Consequently, in the MAR and median grainsize record of Lej da Champfèr the green rhythmites of subunit A are interpreted as sediment deposited during glacier advances and/or periods with the most intensive catchment glaciation and the lowest basal shear stresses. They thus represent the coldest and most continental climate conditions on record. In contrast the grey silts of subunit B are interpreted as the interval with the highest basal shear stresses or the strongest Late Glacial glacier retreat, since it shows high MAR and a broad grainsize maximum. They therefore would represent the warmest period on record. In the time interval between the two continental phases of the green rhythmites glaciers changed several times between the two modes and/or fluctuated between their maximum and slightly retreated positions as indicated by short term, high Late Glacial in Lake Champfèr 146 Chapter 7 amplitude oscillations in the median and MAR values. Interpreted in the sense of Maisch and Haeberli (1982) this would indicate an unstable climate that oscillates between cold (and probably arid) and warmer (and probably humid) phases. There is indication that glaciers might have retreated back to subunit B positions during short intervals (cf. mineralogy section). However, with a moderate retreat and increased mass balance gradient probably also reworking of subglacial till played a more important role for sediment accumulation. For example, the transition from the upper green rhythmites to the grey silts in subunit A is documented as a relatively abrupt shift in the MAR and grainsize records indicating a change in glacier movement most likely an initially fast retreat of the glaciers. The grey silts thus certainly were deposited during a phase of strong glacier retreat, since the unit 3/4 boundary was interpreted as the cessation of glacial meltwater input, i.e. glacier free catchment conditions. In agreement with the above proposed model the highest organic carbon accumulation rates (figure 7.7b) are observed in subunit B. This could indicate a higher lake productivity or more allochthonous input of organic matter. The gradual increase and decrease of the organic carbon accumulation rates (figure 7.7b) argue for the latter possibility. Allochthonous input of organic matter is a reflection of the development of vegetation and soils in the catchment. It is known that the establishment of vegetation does not happen immediately after deglaciation but lags behind (an abrupt) climate warming (Walker, 1995). This would explain the observed gradual change. In addition, this lag time might also be the reason why no significant reduction in organic carbon accumulation is observed for the upper green rhythmites of subunit C. This cold interval might have been too short to trigger a significant change in the vegetation cover and thus the allochthonous input of organic matter. The organic carbon accumulation rate hence represents a smoothed climate signal. Subunit B therefore probably represents the warmest interval on record. Glaciers might have retreated almost to LIA positions (cf. mineralogy section). During the deposition of subunit C glaciers were in advanced positions and had a high mass balance gradient. Continental conditions with glacier positions like in subunit A apparently only occurred for a very short time interval as indicated by a small package of green rhythmites in the second half of subunit C. The transition to subunit B shows up in the grainsize and MAR record as a gradual change (figure 7.6a+e). This indicates a relatively slow retreat of the glaciers as response to the warmer conditions of subunit B. The model of Leemann and Niessen (1994b) thus can be modified for Late Glacial conditions by considering the mode of glacier movement. With respect to the two parameters median grainsize and MAR this implies: Chapter 7 147 Late Glacial in Lake Champfèr

• The median grainsize of the sediment deposited in the proglacial lake is not merely determined by the mixing of non-glacial (rainfall derived) and glacial particles. In contrast, grainsize variations in glacier derived debris caused by changes in the mode of glacier movement play a more important role during the Late Glacial. • The MAR is not solely controlled by the degree of catchment glaciation. In contrast in the cold Late Glacial environment the way of glacier movement is of vital importance for MAR. It has to be considered that low basal shear stress, that char- acterizes glaciers in cold, arid continental climate, can lead to reduced glacial erosion capacity and thus to low MAR in spite of intensive catchment glaciation. However, when interpreting the grainsize and MAR records as a documentation of the glacier history it has to be considered that, different from the present situation, the Late Glacial Lej da Champfèr, in addition to the meltwater input from the glaciers in Val Fex - that was filtered by the passage through Lej da Silvaplauna - might have received significant amounts of meltwater from other sources. For instance the Late Glacial advances and meltback of the nearby Vadret Lagrev (figure 7.14) certainly introduced glacier milk into the Silvaplauna-Champfèr system. During Post- glacial times this small glacier was not important as source of glacier milk. The same holds for the glaciers at the western side of the mountain chain ranging from Piz dal Lej Alv via Piz Corvatsch, Piz Murtel and Piz Surlej to Piz Rosatsch (figure 7.14). There during the Postglacial, glaciers were either very small or not present (Maisch, 1992), whereas the situation was different in the Late Glacial (Suter et al., 1982). Con- sequently, reading the information that is recorded in the Late Glacial sediments of Lej da Champfèr would be much more refined when the importance of the different possible sediment sources can be traced through time. One way to do this is to use mineralogical tracers.

7.3.5 Sediment mineralogy as a key to the glacier positions during the Late Glacial

In chapter 5.3.1 it was shown that the mineral composition of the river and lake sediments can be viewed as a fingerprint of the source region of the suspension. In this section the detected variations of the sediment mineral assemblage are interpreted in this sense. The glacier history read from the analysis of the grainsize and MAR records given above is the baseline for this interpretation. In figure 7.14 the positions of Late Glacial terminal moraines as given in Suter et al. (1982) are superimposed on a simplified geological map (PTT, 1960) to illustrate the different possible source regions of suspension (glacier milk). It is stressed that, the following interpretation of changes in the sediment mineral assemblage is the ‘presently most conclusive solution’ and therefore speculative. Late Glacial in Lake Champfèr 148 Chapter 7

1 1 Lej da 2 7 S. Murezzan 9

Lej da 3 11 Champfèr

8 10 5 4 6 12

9 Lej da Silvaplauna

18 13

Lej da Segl 14 9

23 25

19 28 15

26 16

27 20

0 1 2km 24 22 21

Glaciers in 1973 Serpentine 1 = Piz Güglia 15 = Piz dal Lej Alv rich rocks 2 = Vadret Güglia 16 = Il Chapütschin Little Ice Age (1850) 3 = Pass del Güglia and 17 = Chastelets 2 YD moraine walls 18 = Furtschellas Kromer stadial Greenshists 4 = Piz Lagrev 19 = Vadret da Roseg 5 = Vadret Lagrev 20 = Vad. da Tremoggia (Preboreal) Carbonate Egesen stadial 6 = Vadret Tscheppa 21 = Vad. da Fex bearing rocks 7 = Ova da Suvretta 22 = Vad. da Güz (Younger Dryas) 8 = Ova da Valhun 23 = Val Fex Cristalline Daun stadial 9 = En 24 = Vad. da Fedoz basement 10 = Piz Rosatsch 25 = Val Fedoz 11 = Vadret Rosatsch 26 = Piz da la Margna Lakes 12 = Piz Surlej 27 = Vad. Forno Position of core PCH 94/3 13 = Piz Murtel 28 = Val Forno Talc bearing rocks Rivers 14 = Piz Corvatsch

Figure 7.14: Map of glacier extensions during the Late and Postglacial (cf. figure 2.4) superimposed on a simplified geologic map redrawn from Staub (1946). Only talc, amphibole, chlorite and carbonate rich lithologies are shown. The encircled area is the catchment at the western slopes of the Corvatsch-Rosatasch mountain chain (cf. figure 2.1). Glacier movements in this catchment to a large extent control the mineral composition of lake Champfèr sediments (see text for discussion). Chapter 7 149 Late Glacial in Lake Champfèr

The Late Glacial sediments of Lej da Champfèr represent a mixture of glacial rockflour from three different sources. 1. The relatively large glaciers in Val Fex in the southeastern catchment 2. The glaciers at the northern slope of Piz Lagrev in the northwestern catchment and 3. The glaciers at the western slopes of Piz Corvatsch. Whether also meltwaters from the glaciers in Val Suvretta reached Lej da Champfèr, which presently is not the case, is unknown. Even if this was the case, it would be difficult to distinguish the respective rockflour from that derived from the Vadret Lagrev, because both glaciers mainly are eroding granites, granite-gneisses and diorites of the Julier nappe (figures 2.3 and 7.14). The other 3 possible sediment sources listed above should yield rockflour with distinctive mineralogical signatures. 1. The rockflour from the glaciers in Val Fex should mainly consist of quartz, mica, chlorite and plagioclase and minor amounts of amphibole. It is not expected that this composition changes dramatically when the glaciers advance or retreat in Val Fex since they are always eroding the same lithologies (figure 7.14). Only when they advance further than their position during the Egesen stadial (heavy black line in figure 7.14) they would erode mesozoic carbonates which then would result in a significant dolomite and calcite content of the glacier milk. This was e.g. the case during the Daun stadial (stippled line in figure 7.14). 2. The rockflour from Vadret Lagrev should mainly be composed of quartz and plagioclase and minor amounts of orthoclase, chlorite and amphibole. It also should not change when the glacier is moving since always the Julier granites, gneisses and diorites are eroded. Hence it is not possible to infer glacier positions in this catchment from the presence or absence of a ‘marker mineral’ in the sediments of Lej da Champfèr. Nevertheless, a more important rockflour input from this catchment should lead to higher amphibole and plagioclase intensities in the sediments. Furthermore it has to be considered that this is the only catchment where direct meltwater input into Lej da Champfèr would have been possible. 3. In contrast, the composition of rockflour from the glaciers at the western slope of Piz Corvatsch could change dramatically when these glaciers retreat or advance. As described in section 2.2 ophiolites of the Platta nappe interfinger with mesozoic sediments, granites and gneisses in this small area (encircled in figure 7.14). De- pending on the glacier positions the glacier milk from this catchment may contain certain ‘marker minerals’ that when detected in the sediments of Lej da Champfèr, allow to infer glacier positions in this catchment. Late Glacial in Lake Champfèr 150 Chapter 7

For instance when glaciers are approximately at LIA positions they erode granites and gneisses of the Err nappe. Quartz, plagioclase and mica should be the dominant minerals carried into Lej da Champfèr. When these glaciers advance they will first meet mesozoic sediments (mainly carbonates). Dolomite intensities should increase and calcite should appear as a ‘marker mineral’ in the sediments. On a further advance green schists and gabbros of the Platta ophiolites and finally serpentinites and talc schists are encountered. Calcite may disappear due to dilution, quartz intensities should decrease and as a ‘marker’ the chlorite/mica ratio should increase remarkably. Finally the occurrence of talc should mark maximum glacier extensions. This succession is strongly idealized because the different geologic units interfinger and do not strictly follow one after the other. Moreover, at least 4 glaciers have to be considered and each of these meets the respective rocks at a different time. However, taking the above succession as roughly correct, the following hypothesis is formulated: The composition of glacial rockflour that reached the northern basin of Lej da Champfèr during the Late Glacial was modulated in a characteristic way by the glaciers on the western slope of Piz Corvatsch (encircled area in figure 7.14). Hence observed changes in the mineral assemblage of Lej da Champfèr sediments allow to trace the glacier positions in this particular catchment. Increased glacier activity (regardless of glacier positions) in the northwestern (Güglia) catchment might be recognized by higher amphibole and/or plagioclase intensities. In conjunction with the evidence presented in section 7.3.3 the mineralogical record is unravelled as follows. Subunit C was deposited when glaciers were in an advanced position, eroding carbonates as well as greenschists. This is indicated by average dolomite intensities (figure 7.10d), the occasional occurrence of calcite (figure 7.11c) and above average chlorite intensities (figure 7.10a). A maximum glacier extension is indicated by a thin package of green rhythmites with amphiboles and talc (figure 7.11a,b). This is consistent with grainsize (figure 7.6a) and MAR (figure 7.6e) minima detected between 989-995 cm sediment depth. The beginning of subunit B is marked by a rapid decrease of chlorite (figure 7.10a) and a strong increase of dolomite intensities (figure 7.10d). The erosion of carbonates (e.g. at Chastelets, figure 7.14) must have become relatively more important. Probably glaciers retreated beyond the limit were greenschists are exposed. Both shifts are abrupt indicating an initially rapid retreat of the glaciers in the Corvatsch catchment as a response to rapid warming and/or rapid decrease of precipitation. Peak intensities of plagioclase (figure 7.9d) and orthoclase (figure 7.9c) that occur after the dolomite peak might indicate a more important contribution from the Lagrev catchment (figure 7.14) for a short time. After that, chlorite intensities (figure 7.10a) Chapter 7 151 Late Glacial in Lake Champfèr fluctuate but show a long term increase whereas dolomite intensities still remain above average and show a second maximum in the lower half of the subunit B yellow silts (figure 7.10d). This indicates that glaciers remained in retreated positions. The minimal plagioclase intensities (figure 7.9d) that are observed for this time interval might be viewed as evidence for very small glaciers at Piz Lagrev. Accord- ingly it was pointed out above that subunit B, because of its higher MAR and grainsize values, is considered as the warmest interval on record. Afterwards, glacier advances are indicated by decreasing dolomite intensities. From the deposition of relatively coarse grained sediments with high intensities of orthoclase, plagioclase, quartz and amphibole, it seems, that initially the glaciers in the western catchment (Vadret Lagrev, Vadret Tscheppa) started to advance. Gla- cial meltwaters might have drained directly into the northern or the southern basin of Lej da Champfèr. The transition zone between subunit B and A (~15cm to both sides of the boundary) in general is charcterized by large fluctuations in all mineral intensities, which is viewed as evidence for a ‘flickering’ (short term retreats and advances) of glaciers in the different catchments. In the middle of this transitory phase advancing glaciers at Piz Corvatsch are indicated by high chlorite/mica ratios (figure 7.10c) and the persisting occurrence of amphibole (figure 7.11a). Calcite disappears completely from the sediments (figure 7.11c). As mentioned in 7.3.3, the lower green rhythmites of subunit A are interpreted as sediment deposited during the phase of maximum catchment glaciation. i.e. glaciers in a relatively stable advanced positions. After a short stagnation the highest amphibole intensities and the presence of talc indicate maximum glacier extensions in the western as well as in the eastern catchment. Talc especially indicates glaciation of the Furtschellas region. In the area indicated on figure 7.14 significant amounts of talc are outcropping at the surface (Finger et al., 1982). Studies of the moraine record (Suter et al., 1982) revealed that glaciers (Vadret del Chapütschin, Vadret Mortèl and two other glaciers) advanced significantly over these rocks for the last time during the Egesen stadial, but never during the Postglacial (figure 7.14). The analysis of the river sediments (figure 5.1) proved that even today the rivers draining this catchment (especially the Ova da l’Alp, figure 5.1) contain small amounts of talc, serpentine and also above average amounts of chlorite. However, the lowest mica (figure 7.10b) and highest amphibole intensities (figure 7.11a) and the initially high plagioclase intensities (figure 7.9d) in the green rhythmites are viewed as being the result of a contribution from the Lagrev catchment. The sediment of the Ova dal Vallun that is draining this catchment contains only very little mica but amphibole and large amounts of plagioclase (cf. figure 5.1). It has to be Late Glacial in Lake Champfèr 152 Chapter 7 taken into account that amphibole could as well originate from the Corvatsch catchment and therefore is not the ideal ‘marker mineral’ for the Lagrev catchment. Talc and amphibole disappear simultaneously immediately after they reached their maximum intensities (figure 7.11a+b). Hence glaciers either retreated very fast beyond the limits of their occurrence or retreated moderately and their disappearance is the result of a dilution by other minerals. A scenario of glaciers fluctuating in advanced positions seems more likely for the time interval between the two green rhythmites, since talc and amphibole again occur sporadically in the sediments whereas dolomite intensities (figure 7.10d) remain relatively constant on a low level. It therefore seems unlikely that glaciers retreated as far back as in subunit B. Glacier positions during this interval most likely were comparable to subunit C, except a higher variability, i.e. larger fluctuations have to be assumed. The following glacier advance represented by the upper green rhythmites resembles similar mineralogical features than the first one, except also orthoclase intensities (figure 7.9c) are very low whereas chlorite and dolomite intensities (figure 7.10a+d) are fluctuating like in the initial phase of the first advance. This may be viewed as a less important contribution from Vadret Lagrev. After that in the upper grey silts of subunit A, a relatively fast glacier retreat is indicated by the occurrence of calcite and a large drop in chlorite intensities (figure 7.10a). High plagioclase, orthoclase and quartz intensities (figure 7.9b-d) might indicate retreating glaciers in the Lagrev catchment. The minimum chlorite intensities and the absence of calcite, after the initial peak, might indicate the deglaciation of the Corvatsch catchment. Gradually increasing chlorite intensities and chlorite/mica ratios (figure 7.10c) afterwards might already mirror the transition towards glacier free (rainfall dominated), Postglacial catchment conditions were chlorite/mica ratios usually are >1.5 (data not shown for Lej da Champfèr, cf. figure 5.7 for Lej da Silvaplauna data). The questions that arise are, how does this glacier and climate history that was inferred from the physical and mineralogical sediment properties compare with the local moraine records and how does it fit into the commonly used chronostratigraphic scheme ?

7.3.6 The dating problem

Usually, for lacustrine sediments there is a variety of techniques available that allow to establish an absolute chronology. However, for siliciclastic sediments, especially Late Glacial sediments, absolute dating is a very difficult task for different Chapter 7 153 Late Glacial in Lake Champfèr reasons. First of all the scarcity of organic matter makes 14C-dating extremely difficult even when AMS techniques are used. Radiocarbon dates of samples containing little carbon are often in error (Olsson, 1979). For the same reason and because carbonaceous shell material is also not present, amino-acid racemization as a possible dating method has to be abandoned too (cf. Lister, 1985) Secondly, a palaeomagnetic mastercurve for the Late Glacial of western Europe is not yet well established, as it is for the Holocene (Creer and Tucholka, 1983). Thirdly, it has been shown by e.g. Lister (1985), that thermoluminescence dating yields far too old ages for glacio-lacustrine sediments because the exposure time in meltwater streams is too short to empty the electron traps. In spite of this difficulties an attempt was made to establish a chronology for the 650cm of Late Glacial sediments in Lej da Champfèr. This preliminary chronology is based on a combination of layer counts and AMS 14C-dating. The quality of this chronology is discussed below. A possible source of error for all AMS dates is the fact that turbidites were sampled. This was done because the ‘normal’ sediment contains by far to little organic matter for dating, whereas turbidites usually are enriched in terrigenous plant remains. The drawback of course is that the dated plant remains may have been re-deposited with the turbidite. This would yield ages that are too old. However, it is difficult to estimate for how long the dated plant remains (needles and twigs) remained in the delta sediments or wherever they have been deposited originally, before they were transported to the profundal sediments. Therefore error estimates can not be given here. In view of these possible processes it seems reasonable to assume that the determined dates rather are too old than too young. The measured ages therefore are considered to represent maximum ages. The same assumption is made for the layer counts. It seems rather unlikely that layers were missed. In contrary, it is possible that too many layers were counted because it was not always possible to distinguish unequivocally between turbidites and ‘varves’. Therefore, I assume that rather too many than too view layer couplets were counted. Concerning the bulk AMS-date of 17970 BP that was rejected because of a possible contamination with old carbon (cf. 7.2.3) it is mentioned here that other dates measured on bulk organic matter exist from the Upper Engadine that are also significantly older than 13000 BP (Ariztegui, 1993; Heitz et al., 1982a; Heitz et al., 1982b; Kleiber, 1974). They all were rejected by the respective authors in view of the above mentioned problem or due to a mismatch with the pollen stratigraphy. However, until 14C-dates on terrestrial macro fossils are available, the possibility of an early Late Glacial in Lake Champfèr 154 Chapter 7 ice-free Upper Engadine perhaps already during the Lascaux interstadial (17000- 19000 BP, cf. figure 2.5) as discussed by Kleiber (1974) can not be ruled out completely. There are 3 possible ways to construct a preliminary chronology for the Late Gla- cial sediments. 1. The AMS date of 9350 BP (10299±255 cal. BP) is chosen as anchor point for layer counting 2. The date of 10000 BP is chosen as the anchor point for layer counting 3. A chronology is constructed that extrapolates from the two most reliable AMS dates Arguments in favour and against each of these chronologies are presented below. 1. Reasons in favour of the correctness of the 9350 BP AMS date are: (1) It was determined on identified terrigenous macro remains (needle remains, small twigs, a woody fragment and cat scales). (2) It is supported by a similar date (9415 ±90 BP) also measured on terrigenous macro remains from the same lithologic boundary in adjacent Lej da Silvaplauna (Leemann and Niessen, 1994b and figure 7.13). From the anchor point at 10299 cal. BP layer counts are used to construct the chronology assuming an annual nature of sediment deposition for the 650 m of pre 10299 cal. BP sediments. The resulting chronology is shown in figure 7.12b as the heavy black line. The only other AMS date that plots exactly on this line is the stratigraphically lowest one at 950 cm sediment depth with 12199 cal. BP. This can be viewed as evidence in favour of the true annual character of the Late Glacial layer couplets. All other AMS dates seem to be too old for they might contain re-deposited material (see above). However, the data shown in figure 7.12 also visualizes a discrepancy concerning the definition of the Late Glacial/Postglacial boundary that arises when this dating is true. The timing of this transition is crucial because it is commonly assumed that the drastic environmental change at this boundary (time equivalent to the Younger Dryas/Preboreal or Pleistocene/Holocene boundary) occurred fast and broadly synchronous at 10000 BP throughout the Alps and perhaps globally. For the late Pleistocene and Holocene the chronostratigraphic scheme introduced by Mangerud et al. (1974) is commonly referred to, which places the Late Glacial/ Postglacial boundary at 10000 14C years BP. This absolute chronology, based on radiocarbon dating was tied to the boundaries of pollen zones defined earlier by Firbas (1954) in Scandinavia, which have been proven to be time-transgressive in other Chapter 7 155 Late Glacial in Lake Champfèr parts of Europe (see Ammann and Lotter (1989) for Switzerland). Nevertheless, in sediment cores from Perialpine lakes it is often observed that a distinct change in pollen assemblage occurs at 10000 BP e.g. Burga (1988). This also seems to be true for high alpine pollen records derived from small lakes and ponds (Ammann et al., 1993). In most lakes this change in catchment vegetation is more or less synchronous with a drastic sedimentological change (Lister, 1988; Niessen and Kelts, 1989). This often is a change from dominantly allochthonous to autochthonous sedimentation. As mentioned above, in Lej da Champfèr this abrupt change in sediment composition is caused by the cessation of glacial meltwater input. Fischer (1996) could show in varve and AMS dated sediments of Soppensee (Switzerland) that the pollen change starts at the same time as the actual change in climate (inferred from the oxygen isotope record) but is more gradual. In contrast the lithologic change in Soppensee clearly predates this change. For the high Alps e.g. Tinner et al. (1996) proposed that lithologic and vegetation changes both occur at 10000 BP. However, the fact that a 14C plateau occurs exactly at 10000 BP (Kromer and Becker, 1993; Lotter et al., 1992c) makes any precise dating of this boundary very difficult. In a pollen record from Lago Basso (2250 m a.s.l.) only ~35 km west of Lej da Champfèr Wick and Tinner (1997) could prove by macrofossil analysis that first trees of Betula pubescens and Pinus cembra arrived near the lake around 9700 BP. From the pollen analysis they conclude that the lithologic change from silt to fine-detritus gyttja that occurs below this date represents the beginning of the Holocene at 10000 BP. In the light of these results the possibility has to be considered that the 14C date of 9350 BP for this boundary in Lej da Champfèr might be too young. 2. For this reason as the second possible chronology, in figure 7.12b this boundary tentatively has been set to 10000 BP and layer counts then were anchored at this date. The resulting chronology (thin black line in figure 7.12b) matches 3 of the 4 AMS dates that were previously classified as too old. Also two declination dates (figure 7.12c), i.e. label Q dated by Lister, (1985); Niessen and Kelts, (1989); Wessels, (1995) and label I dated by Wessels (1995) plot on the thin black line and therefore would constrain the assumption of a lithologic change at 10000 BP and not at 9350 BP. Declination swings below I (I-j to I-m) are either labelled incorrectly or indicate significantly lower sediment accumulation rates prior to 11000 BP. 3. The third possible chronology is represented in figure 7.12b by the heavy grey line. This chronology exclusively relies on the two AMS dates at 505 and 639 cm sediment depth that were determined on terrestrial macrofossils (table 7.1). The dates prior to 11492 cal. BP were derived by linear extrapolation assuming constant accu- Late Glacial in Lake Champfèr 156 Chapter 7 mulation rates. The AMS date at 12257 cal. BP also plots on this line, whereas the others are too old or too young. The four lowermost declination dates (label I-j to I- m in figure 7.12c) also would agree with this chronology. In contrast this chronology would imply that the layer counts are not correct. With the information available at the moment no unambiguous decision can be made whether one of the proposed chronologies is definitely correct. Also the first results from pollen analysis that started at the beginning of this year are not yet conclusive (E. Gobet, pers. com.). The only way to solve this dating problem in the future is to establish a high resolution pollen stratigraphy in combination with high resolution AMS dating for the Late Glacial sequence and especially across the unit 3/4 boundary. This is one topic that will be addressed in the PhD-thesis of E. Gobet (in prep.). However, it is remarkable that for either of the two proposed chronologies at least between 505 and 950 cm sediment depth an agreement can be found between the layer counting and AMS dating. This might indicate that counted couplets in this interval indeed are true varves. For the purpose of this study the counting of layer couplets is favoured as the currently most reliable chronologic tool. As mentioned above it seems rather unlikely that to few layers were counted. Furthermore, the prerequisites for varve formation, i.e. the seasonally varying input of suspension and the summertime thermal stratification of the lake most likely also existed during the Late Glacial. If deposition would be non-annual it is much more likely that a layer couplet represents one of perhaps several meltwater events in one year and not one meltwater event in a couple of years. Again the error then would be too many rather than too view counted years. Therefore layer couplets tentatively are considered as true varves. The layer counting is regarded as a floating varve chronology.

7.3.7 Climatic interpretation and comparison with other records

When we accept the above interpretations of (1) the green rhythmites as representing the largest catchment glaciation during times of cold and arid climate and (2) the counted layers being annual in nature, the duration of the cold phases can be estimated. From the lower to the upper green rhythmites we get 35, 450 and 170 years for the durations of maximum glacier extensions. Based on the climatic interpretation given under 7.3.4 and 7.3.5 chronozones have been assigned to the sediment record as shown in figure 7.15. The median grainsize curve is plotted as representative for the glacier- and thus climate history in the way Chapter 7 157 Late Glacial in Lake Champfèr described above. The terms Younger Dryas (YD), Allerød (AL) and Bølling (BO) are used merely in a climatic sense. The varve chronology may be confirmed based on this interpretation.

Geomorphological background In geomorphology glacier advances usually are expressed as ELA-depressions. This is the lowering of the equilibrium line altitude (ELA) of a glacier with respect to a well known reference level (often ELA at AD 1850). From the Late Glacial terminal moraines palaeo-glacier geometries can be reconstructed, which then are used for ELA calculations assuming a 2:1 ratio of accumulation to ablation area (Gross et al., 1976). An ELA-depression is the expression of a positive mass balance of a glacier which in turn can be the result of both increased precipitation or reduced temperature. For instance for the AD 1850 LIA glacier highstand in the Lej da Silvaplauna catchment Maisch (1992) calculated an ELA depression of 130 m with respect to the 1973 reference ELA. This can be produced by a lowering of average annual air temperature by 1°C or an increase in annual precipitation of 500 mm or a combination of both.

The Egesen (YD) ELA depression In the Upper Engadine glaciers advanced significantly beyond their Postglacial (LIA) maximum positions for the last time during the Egesen stadial (Alp Val in local terminology, figure 2.5). As shown in figures 2.4 and 7.14 this glacier highstand is well recorded by several terminal moraines. E.g. for the Vadret da Roseg, 4 glacier advances belong to the Egesen stadial (figure 2.4). In the catchment of the Upper Engadine lakes most times 2 Egesen moraines are recognized (figure 7.14), e.g. Val Fex (Suter et al., 1982). Furrer et al. (1987) note an ELA depression of 170-245 m compared to the 1850 reference level for the Egesen stadial. This indicates mean annual air temperatures 3-4 °C lower than today (Haeberli, 1991), which agrees with estimates made e.g. by Eicher and Siegenthaler (1976) from d18O-measurements on lake carbonates and Kerschner (1980) made from ELA depressions, timberline altitudes and the distribution of fossil rock glaciers.

The YD cold interval The Egesen stadial is equivalent to the Younger Dryas (YD) chronozone, a climate deterioration that affected the north Atlantic realm, all areas of Europe (Walker, 1995) and perhaps is recognized globally at the same time (e.g. Ariztegui et al., 1997; Peteet, Late Glacial in Lake Champfèr 158 Chapter 7

1995). The length of the YD meanwhile in most terrestrial and marine records is approaching the length inferred from several Greenland icecore studies, which is 1000 - 1200 calendar years (e.g. Björck et al., 1996; Goslar et al., 1995; Hajdas et al., 1993; Johnson et al., 1992). In the icecores and in different archives around the North Atlantic the YD cold interval occurs synchronously between 12600 and 11400 cal. BP. However, there is still considerable scatter in the absolute dates for the YD boundaries in other archives. Hence, the world-wide synchroneity is still debated as are the possible forcing mechanisms. E.g. a global synchroneity would demand a coupling between ocean circulation changes and atmospheric circulation. There is evidence from several different archives that climate amelioration took place during the second half of the YD (Goslar et al., 1993). Probably climate changed from cool and humid to drier, more continental (see compilation in Wohlfarth et al., 1994). Another YD-type cold event that is also recognized in many archives around the North Atlantic occurs 300 to 500 years after the termination of the YD during the Preboreal (see compilation in Behre, 1978; Walker, 1995). This Preboreal-oscillation is shorter (150 years in icecores, Björck et al., 1996) and less severe than the YD. It is e.g. also recognized as glacier advances in Sweden (Nesje and Dahl, 1993). However, in the Upper Engadine moraine record this oscillation is not well documented (figure 7.14). It could be equivalent to the local Beverin or Ova d’Err moraines (Kromer or Schlaten-stadial in the Eastern Alps, figure 2.5). With 90-120 and 60-70 m the ELA- depression of this glacier advance is clearly smaller than for the Egesen stadial.

Green rhythmites = YD With this information it is most likely that the lower green rhythmites of subunit 4A were deposited during the YD glacier advances. But with 450 years duration their time interval of deposition is not even half as long as the length of the YD. There are three possible reasons for this. (1) The green rhythmites represent only a part of the YD, (2) Layer couplets are not annual and accumulation rates according to chronology 3 proposed under 7.3.5 are 2 to 2.5 times lower than given in figure 7.6e, (3) The above hypothesis of low median grainsize representing glacier highstands is incorrect. The latter seems rather unlikely since other studies constrain this hypothesis (Ariztegui et al., 1997; Leemann and Niessen, 1994b) and it was demonstrated above that it is consistent with other lines of evidence in this study. Point (2) would imply an age for the core base of 17100 cal. BP (14500 BP). In this case one should expect that the much larger glacier advances of the Palüd Marscha, Bever and the Chinuos- chel stadial (Daun and ~Geschnitz of the Eastern Alps, figure 7.14) with ELA-de- Chapter 7 159 Late Glacial in Lake Champfèr pressions of 240-250 and ~500 m, respectively which occurred during the older Dryas or the oldest Dryas, would be recorded in the sediments as well. There is no evidence for this, which would be green rhythmites like sections of reduced median and different mineralogy. Furthermore, the average Late Glacial sedimentation rates then would be in the same order of magnitude than the Neoglacial rates (0.1 cm/yr). In view of a much larger catchment glaciation and a sparser vegetation cover during the Late Glacial this seems rather unlikely. Hence, if we consider point (1) as being correct two questions remain: 1. Which part of the YD do the lower green rhythmites represent and do the upper green rhythmites also belong to the YD glacier advance ? 2. How does the glacier history deduced from the sediment record compare to the glacier history that was reconstructed from the moraine record (summarized in Suter et al., 1982) ?

Younger Dryas For the moment an interpretation that regards the two packages of green rhythmites of subunit 4A as the beginning and the end of a cold interval equivalent to the YD seems the most logical assumption for the following reasons. The shift in grainsize and mineralogy that is observed at the base of subunit 4A is very abrupt. This indicates a fast glacier advance were glaciers ‘suddenly’ eroded different lithologies as a response to a rapid temperature drop as e.g. evident from the Greenland icecore d18O-record (Dansgaard et al., 1989). Afterwards constant mineral composition and grainsize as well as unprecedented rhythmic sedimentation indicate that glaciers remained at their maximum positions for 450 years. This constrains the findings of Ivy-Ochs et al. (1996) who propose exactly the same (rapid advance followed by a long maximum) based on moraine morphologic features of the outer moraine wall at Pass del Güglia (figure 7.14). The inner moraine at Pass del Güglia according to Ivy-Ochs et al. (1996) rather shows features typical for ice rich permafrost. They, in accordance with other records (cf. Wohlfarth et al., 1994), propose a formation under drier, more continental conditions with starving glaciers during the latest phase of the YD. Different from that the sediment record suggests that after a phase of glacier fluctuation and probably more unstable climate (cf. 7.3.8) a second glacier advance represented by the upper green rhythmites occurred that was as severe as the first YD advance. A similar magnitude is proposed because the upper green rhythmites show the same grainsize and mineralogical signatures as the lower green rhythmites. This second advance then started 350 years after the first one ended and lasted only 150 Late Glacial in Lake Champfèr 160 Chapter 7 years. On the other, hand there is mineralogical evidence that input from the Vadret da Güglia was reduced at that time, which again would agree with Ivy-Ochs et al. (1996) suggestion of a rock glacier at Piz Güglia rather than a ‘real’ glacier with YD dimensions at that time. For the two moraines at Pass del Güglia (figure 7.14) Ivy- Ochs et al. (1996) determined exposure ages of 11800 cal. BP for the outer and 10400 cal. BP for the inner wall. Within the limits of error the first date would fit well with the first chronology proposed in 7.3.5 (figure 7.15). The second date would be younger than the advance represented by the upper green rhythmites which then is centred at ~10900 cal. BP. This date would thus plot into the Preboreal. As mentioned above all Preboreal glacier advances show ELA-depressions that are significantly smaller than those determined for the YD. It is therefore rather unlikely that the upper green rhythmites represent a Preboreal glacier advance. Most of the other Egesen moraines that were detected in the southeastern catchment also show two walls. These usually are only a few meters apart and thus represent approximately the same ELA-depression. Therefore it is much more logical to assume that the lower and the upper green rhythmites, represent two glacier advances within the Egesen stadial. Whether the transition to subunit A grey silts or the unit 3/4 boundary is representing the YD/PB transition remains unclear. The length of the YD would be 1000 years in the first and 1250 years in the latter case. Until new palynological evidence is available it is suggested that the boundary between the upper green rhythmites and the grey silts of subunit 4A marks the end of the YD. It is hence proposed that glaciers in the southeastern catchment advanced again to their maximum positons and remained there for 150 years at the end of the YD. Respecting the hypothesis formulated under 7.3.4 (low MAR and median as indicators for low glacier basal shear stresses as a result of cold arid climate) it is concluded that the lower and upper green rhythmites represent cold continental phases at the beginning and the end of the YD. In the 350 years in between these continental phases climate probably was more instable. Glaciers during this interval fluctuated in advanced positions, probably due to larger variations in temperature and/or precipitation. This interpretation does not agree with other records, that show humid conditions at the beginning and an increase in continentality in the second part of the YD (Wohlfarth et al., 1994).

Glacier retreat rates at the YD/PB transition Based on the mineralogical evidence (7.3.4) it is further proposed that glaciers retreated very fast from these maximum positions of the latest YD. Taking the mini- Chapter 7 161 Late Glacial in Lake Champfèr Cc O 18 δ (1996) Fischer -8 -7 -6 -5 [‰ PDB] -9 -10 650 1150 Soppensee carbonates 10986 12150 12800 cal. yrs. BP O 18 δ [‰] (1993) Dansgaard GRIP-icecore -43 -41 -39 -37 -35 BP cal. yrs. 11100 11300 11500 11700 11900 12100 12300 12500 12700 12900 13100 13300 13500 13700 13900 14100 14300 14500 5pt RAv [µm] median Grainsize 3.0 4.0 5.0 6.0 946 647 This study cal. yrs. BP 10304-10814 11250-11760 11897-12407 Varve chronology anchored at 10299 cal. BP PB AL YD BO C A B 4 4 4 5 6 7 8 9 10 11 [m] Sed. depth

Figure 7.15: Chronology for the Late Glacial sediments of PCH 94/3 based on layer counting and climatic interpretation of the sediment record. Note the good agreement for the length of the YD and the AL chronozones with the Soppensee oxygen isotope record (Fischer, 1996). Dashed grey lines are proposed correlations with the Greenland isotope record from Dansgaard (1993). Late Glacial in Lake Champfèr 162 Chapter 7 mum chlorite intensities and the occurrence of calcite (cf. 7.3.4) as evidence that glaciers retreated at least back to their LIA positions, average retreat rates can be calculated for the glaciers at the western slope of Piz Corvatsch (figure 7.14 encircled area). There the glaciers must have retreated around 1000 m in 30 to 50 years. The average retreat rate thus is 20-33 m/yr. Assuming that the Vadret da Fex (figure 7.14) also retreated to LIA positions. i.e. a retreat of 3500 m in 30 to 50 years, the average retreat rate for this larger glacier amounts to 70-120 m/yr. For comparison the same glacier (Vadret da Fex) retreated by 1600 m from its LIA maximum in 1850 to the position in 1973. This represents an average retreat rate of 13 m/yr and thus is an order of magnitude smaller than the proposed Late Glacial retreat rates, even if it is taken into account that usually larger glaciers retreat at higher rates. Assuming the interpretation of mineralogical tracers is correct the unprecedented high retreat rates support the idea, that deglaciation also in the central Alps occurred as rapid ice decay rather than a gradual retreat of valley glaciers. This illustrates that compared to the Postglacial, glacier variations the Late Glacial environmental changes really were catastrophic events. Hence the idea of a very rapid climate warming at the YD/BP boundary as was suggested by numerous authors (e.g. Dansgaard et al., 1989; Fischer, 1996; Goslar et al., 1993) is supported by these findings. The grey silts of subunit 4A would then represent the early part of the Preboreal were glaciers retreated at Postglacial rates (14 m/yr) until they vanished completely at the unit 3/ 4 boundary. During these 200 years minor fluctuations might have occurred. But the mineralogical tracers suggest that glaciers most likely never extended significantly further than approximately LIA positions.

Allerød When we accept the base of subunit 4A as the beginning of the YD, subunit 4B would then represent the Allerød interstadial. In the numerous pollenstratigraphies from the Swiss plateau this period usually is seen as the establishment of a pine- birch woodland. In the annually laminated sediments of Soppensee (Fischer, 1996; Hajdas et al., 1993) the Allerød lasted 650 cal. years (figure 7.15). In Lej da Champfèr subunit 4B has a duration of 647 years (figure 7.15). It was shown above that the grainsize, the organic carbon accumulation and the mineralogical evidence consistently indicate that subunit 4B represents the warmest interval in the Late Glacial sediment record. Glaciers probably retreated back to LIA positions in some periods. Pollenstratigraphies and the Gerzensee isotope stratigraphy (Eicher and Siegenthaler, 1976) show a short oscillation in the late Allerød (Gerzensee fluctuation, Lotter et al., 1992a), approximately 200 years before the beginning of the YD. From the Lej da Chapter 7 163 Late Glacial in Lake Champfèr

Champfèr record it is evident that unlike the Aegelsee fluctuation (see below) this climate oscillation did not cause glacier advances to YD limits (figure 7.15).

Bølling Prior to subunit B the first 750 years of the Lej da Champfèr record (subunit 4C) must then represent the last part of the Bølling interstadial (figure 7.15). In pollenstratigraphy the Bølling usually is recognized as the reforestation of the Swiss plateau with birch (Ammann et al., 1994). The Lej da Champfèr sediment proxies indicate intermediate glacier positions, i.e. smaller than YD but larger than LIA, for this interval. Towards the end of subunit 4C a small package of green rhythmites indicates that YD-type glacier dimensions might have been reached for a very short time period of 35 years (figure 7.15 and 7.10c). The latter would match well with pollenstratigraphies that often also show a short regressive phase in vegetation development (e.g. in Soppensee) commonly referred to as the Aegelsee fluctuation (Lotter et al., 1992a). Morphological witnesses of this advance were then destroyed by the following, stronger YD advances. However, considering the oxygen isotope records from Gerzensee (Eicher and Siegenthaler, 1976) or the GRIP icecore (Dansgaard et al., 1993) (figure 7.15) the Bølling should be marked by a strong initial warming, with temperatures higher than during the Allerød. This is either not recorded in the sediments or was not captured in the core because it is too short. When this climatic interpretation and the assignment of chronozones is correct the resulting absolute chronology would be in good agreement with the relatively well constrained varve chronology from Soppensee (Fischer, 1996; Hajdas et al., 1993; Lotter et al., 1992b). This would again demonstrate the discrepancy between radiocarbon and sidereal years and the existence of an offset with respect to the Greenland icecore chronology which is observed in many terrestrial records (figure 7.15). For comparison from the absolute chronologies proposed under 7.3.5 the one that is anchored at 9350 BP is also given in figure 7.15. It is seen that also within the limits of error this chronology would be younger than e.g. the Soppensee chronology that is constrained by extensive AMS dating. The median grainsize curve, if taken as palaeoclimatic indicator in figure 7.15 resembles roughly similar trends than the other two high resolution records. For the ice core and the lake carbonate records a climate control of the oxygen isotope ratios is not questioned any more. Therefore, from a phenomenological point of view it seems likely that all 3 the archives recorded the consequences of the same climatic event. Late Glacial in Lake Champfèr 164 Chapter 7

7.3.8 Comparison of interannual variability in different MAR records

It was shown in 7.3.4 that the model of Leemann and Niessen (1994b) concerning the relationship between MAR and glacier size had to be modified for the Late Gla- cial sediments of Lej da Champfèr. Therefore, a direct comparison of recent MAR from Lej da Silvaplauna (Leemann, 1993) with Late Glacial MAR from Lej da Champfèr on decadal to centennial time scales should not be attempted. An approach that might be feasible is to compare interannual variability of MAR as observed in the calibration period (chapter 4) with interannual MAR variability during selected Late Glacial time intervals (figure 7.6e). Figure 7.16 shows this comparison for 4 normalized 130 year time periods taken from different parts of the varve records. These are the Lej da Silvaplauna calibration period (AD 1864-1993,

6 figure 7.16a) and from MAR (AD 1864-1993) 5 Lej da Champfèr the 4 3 Allerød, the first and the 2 1 second half of the YD 0 (figure 7.16b-d). Look- -1 ing at the Late Glacial 5 MAR (YD second half) 4 varves, on the first 3 glance there is no funda- 2 1 mental difference in 0 -1 interannual variability. This indicates that no 5 MAR (YD first half) 4 matter if climate is in a 3 warm or in a cold 2 1 ‘mode’, meteorological

MAR (normalized values) 0 -1 conditions always change in the relatively 5 MAR (Allerød) 4 narrow range that we 3 2 experienced in historical 1 times. The biggest dif- 0 -1 ference between the -2 0 10 20 30 40 50 60 70 80 90 100 110 120 130 records shown in figure Varve years 7.16 is the fact that the Figure 7.16: Comparison of 130 year long intervals of MAR from a) the record from the second recent (AD 1864-1993) sediments of lake Silvaplana (PSP 94/3-0), on the half of the YD, which one hand with sediments from lake Champfèr (PCH 94/4) from b) the second half of the YD, c) the first half of hte YD and d) the Allerød on the other was interpreted as an side. All values are normalized (cf. 3.9). Note the more frequent occurence interval with larger gla- of ‘high MAR-events’ in the second half of the YD. Chapter 7 165 Late Glacial in Lake Champfèr cier fluctuation based on grainsize and mineralogical data (figure 7.6a and e.g. 7.10c) apparently is punctuated by several events of unusually high MAR. This might be due to, glacier specific processes or might indeed reflect a more instable climate with larger interannual oscillations in temperature and/or precipitation at that time.

7.4 Conclusions

The physical and chemical sediment properties and the mineral assemblage of the 1159cm core PCH 94/3 from the northern basin of Lej da Champfèr exhibit distinct changes in the Postglacial as well as in the Late Glacial part of the core. The most drastic, abrupt shift occurs at the unit 3/4 boundary. This shift is detected in all 4 Upper Engadine lakes and most likely marks the cessation of glacial meltwater input. It is suggested that this cessation did not happen simultaneously in all lakes. Probably Lej da S. Murezzan received meltwaters 500 years longer then the two upstream lakes. Highest susceptibilities in Lej da Segl and Silvaplauna Late Glacial sediments suggest input of magnetic minerals from glacial erosion of ophiolites in the southeastern catchments. Constantly low susceptibilities indicate that direct meltwater input from these catchments never occurred in Lej da Champfèr and S. Murezzan and stopped at the unit 3/4 boundary for the other 2 lakes. In contrast at this boundary for Lej da Champfèr the cessation of carbonate input points to direct meltwater input from the southeastern catchment during the Late Glacial. Appar- ently Neoglaciation was not strong enough to activate this sediment source again. It was demonstrated that the physical properties of Lej da Champfèr sediments show distinct changes during the Late Glacial. These changes most times are accompanied by changes in the mineral assemblage. Taken together these sediment parameters can yield valuable information about the glacier history in the lake’s catchment. It is suggested that the median grainsize reflects the catchment glaciation whereas the mineralogy allows to trace the position of glaciers in the different catchments. Based on their small median grainsize and their distinctive mineralogy the green rhythmites have been identified as a sediment type that was deposited during the most extensive catchment glaciation. A model is suggested that explains the coinciding low MAR with lower basal shear stresses of larger and colder Late Glacial glaciers. According to this model, the green rhythmites would represent times of cold, dry continental climate. Absolute dating of the Late Glacial sediments due to the scarcity of organic matter is ambiguous yet. Therefore, tentatively an absolute chronology is proposed that Late Glacial in Lake Champfèr 166 Chapter 7 relies on the climatic interpretation of the sedimentary proxy indicators and the annual nature of the counted layer couplets. This process-oriented dating approach may be viewed as an ‘educated guess’. It is stressed that interpretations remain speculative until more AMS dates and a detailed pollenstratigraphy are available. Espe- cially the latter would be important because it can provide independent evidence in favour or against the climatic interpretation of the sediment properties. According to the chronology proposed in figure 7.15 core PCH 94/3 would reach back into the Bølling. Sediment median and mineralogy would then indicate glacier positions between LIA and YD for the Bølling. A short advance to YD-type extensions, probably time equivalent to the Aegelsee oscillation is indicated towards the end of the Bølling. The Allerød is regarded as the warmest period with glaciers probably back at LIA positions. Glaciers then advanced in response to the temperature decrease and/or precipitation increase at the beginning of the YD. They reached their maximum positions in less than 100 years. For the first 450 years of the 1000 year long YD, glaciers remained at their maximum positions. The following 350 years are characterized by glaciers fluctuating around their maximum positions, indicating climate conditions that are more instable. Larger grainsize variations and the comparison of YD-MAR with MAR in other time intervals suggests that interannual and interdecadal variability of temperature and precipitation probably was higher. During the last 150 years of the YD a second glacier advance occurred that was as severe as the first one, indicating colder climate again. From these maximum positions glaciers retreated initially very fast probably at a rate of 70-120 m/yr during the first 30-50 years. The following complete meltback of glaciers during the early PB then occurred with retreat rates that were similar to rates determined for the Postglacial (14 m/yr). Thus from the sediment record it seems that the YD was not a uniform cold period. Apparently considerable fluctuations occurred during its second half. Already Kerschner (1980) pointed out that the different YD ELA-depressions of north facing glaciers on the northern slopes of the Alps and in the central area are a result of differences in precipitation. Even if the above interpretations will be modified in future, the fact that abrupt shifts of physical properties and mineralogy were detected in the Late Glacial sediments demonstrates that these sediment signatures can be a powerful tool in reconstructing Late Glacial positions of glaciers. Depending on catchment geology it may even be possible to infer rates of glacier size change. Hence, the laminated sediments of Lej da Champfèr can be viewed as a high resolution archive where glacier and thus climate history was recorded continuously but in a complex way. Therefore, Chapter 7 167 Late Glacial in Lake Champfèr similar to the pioneer studies in isotope stratigraphy of lacustrine carbonate precipitates (e.g. Eicher and Siegenthaler, 1976) a better understanding of the processes controlling the formation of the proposed proxy indicators has to be achieved in the future. In particular a key issue certainly is the validation of these proxies with palynological evidence that is currently under way. Moreover, it is necessary to understand the grainsize and mineralogical characteristics of the glacier milk for different lithologies and different glacier characteristics. Owing to the absence of authigenic carbonate and the scarcity of organic matter it seems unlikely that it will be possible to use isotopes as palaeoclimatic proxy indicators in proglacial lacustrine sediments. A better understanding of the alternative proxies presented in this study and the study of Leemann (1993) therefore is an important task.

Chapter 8 169 Conclusions

CHAPTER 8

Conclusions and Outlook

8.1 The use of siliciclastic sediments as chronicles of climate

The ultimate goal of this study was to compare the annually laminated clastic sediments of Lej da Silvaplauna with instrumental meteorological data with the aim to obtain a transfer function which in a second step can be applied to the Late Glacial sediments. The anticipated translation of annual mass accumulation rates (MAR) into meteorological data then would allow the quantitative interpretation of the drastic Late Glacial climatic changes. The results presented in this thesis show that this is complicated for the following reasons: • Interannual MAR changes are controlled by local meteorological conditions in a complex way. At least three different climatic forcing factors are important. Their influence on MAR shows a high temporal variability (chapter 4). • These multiple controls on varve formation, can not be traced separately by means of geochemical, mineralogicalor magnetic tracers. The importance of early human impact is unclear yet (chapter 5). • Only MAR anomalies reflect meteorological extreme conditions of regional significance (chapter 4+6). • The response of the MAR signal to climatic forcing presumably is non-linear (chapter 6). • Long-term MAR trends in the Late Glacial sediments of Lej da Champfèr do not agree with the model proposed for the Postglacial sediments of Lej da Silvaplauna, which links higher MAR to larger catchment glaciation (chapter 7). Instead they reveal that the mode of glacier movement plays a major role in the Late Glacial. These 5 points caution against the application of linear transfer functions in paleoclimatic reconstructions from proglacial varves. Especially the last point questions the validity of an actualistic approach in sediment studies, because it clearly shows that for the Late Glacial new parameters come into play that were not predict- able from an analysis of the recent sediments. Based on these findings the results of this thesis can be summarized as follows: Conclusions 170 Chapter 8

One important result is that it could be demonstrated that variations in MAR can not be explained by variations in mean summer air temperature alone. It was shown that at least one MAR anomaly in the 130 year calibration period was caused by increased summer rainfall and thus is not dominated by temperature. Statistical analysis suggests that there is a probability of ~50% that a given fluctuation in MAR is caused by a summer temperature anomaly. However, disregarding the cause of an anomaly it could be demonstrated that positive as well as negative MAR anomalies almost always reflect meteorological extreme conditions which are detectable on a regional scale and perhaps linked to large scale atmospheric circulation of western Europe. The futile search for a method that allows to quantify the contribution of rainfall to annual MAR lead to the second important result. Not surprisingly it was found that the mineral assemblage in the sediments of the tributary rivers reflects quite well the geology of the catchment they are draining. Turbidite mineralogy suggests that this mineralogical fingerprint is transferred to the lake sediments. Because the lake catchment is geologically diverse it was possible to use sediment mineralogy to trace the positions and rates of movement of Late Glacial catchment glaciers. In conjunction the sediment median grainsize in accordance with existing models is viewed as a tracer for the degree of catchment glaciation. As mentioned above modification of the existing model is necessary to explain Late Glacial MAR variations. Long- term MAR minima in this time interval occur during maximum catchment glaciation, whereas in the Neoglacial record MAR minima represent smaller glaciers. This fundamental difference is the result of altered boundary conditions. Most likely it reflects the different way of glacier movement of the larger Late Glacial glaciers. Late Glacial changes in MAR do not only reflect the size of the catchment glaciers. This is true for the temperate Postglacial glaciers. In contrast Late Glacial glaciers, although larger, most likely were not temperate at all times. Therefore, a model was proposed where lacustrine MAR are controlled by the basal shear stress of the glaciers. In the Late Glacial according to Maisch and Haeberli (1982) under cold, arid continental conditions cold glaciers with low basal shear stress prevailed at certain times. Low basal shear stress in turn leads to low erosion capacity and thus low MAR in the receptive proglacial lake. It was thus concluded that low MAR coinciding with low median in the Late Glacial sediment of Lej da Champfèr reflect periods of cold and arid continental climate. Similarly a MAR minimum detected in the middle of the LIA between AD 1738 and 1766 might also be the result of this phenomenon. Assuming an annual nature of the Late Glacial layer couplets a third important result is the rapidity of glacier size changes reflected by abrupt changes in median grainsize, MAR and mineralogical composition of the sediments. With all restric- Chapter 8 171 Conclusions tions mentioned in chapter 7 it seems that glaciers retreated from their YD maximum positions within 30-50 years at a rate of 70-120 m/yr. This is a 5 to 10-times higher rate than the rate of 13 m/yr that was determined for the glacier retreat that followed the LIA maximum. Moreover, all tracers indicate that the YD was not a uniform cold interval. According to the model suggested above the first 450 and the last 150 years of the YD are characterized by a cold continental climate. For the 350 years between these two continental phases, in the second half of the YD a more instable climate with fluctuating glaciers is indicated by larger interannual MAR variability.

8.2 Proposed future research

It was demonstrated that the siliciclastic, laminated sediments of the Upper Engadine proglacial lakes record environmental changes in a high resolution mode. To improve the reading of climatic information encoded in the sediments the following suggestions are made. Chapter 4 revealed that the contribution of soil erosion caused by rainfall probably is the biggest uncertainty. Therefore, isolating a tracer that allows to quantify this contribution to annual MAR is essential for a more quantitative interpretation of the varve record. The analyses presented in chapter 5 demonstrated that this is not an easy task. It was suggested that the most promising potential tracer might be found in the magnetic properties of the sediments. To evaluate this a more detailed analysis of the sediment magnetic properties is necessary. This then would have to be validated with e.g. geochemical and biological sediment signatures. In this context it would also be valuable to conduct a sediment trap study. When this is done with adequately high time resolution (every 1 to 2 weeks) it would allow the magnetic, chemical, biological and mineralogical characterization of material deposited under known meteorological and hydrological boundary conditions. Ana- lyzing the suspension load of the catchment rivers in the same time interval would enable us to translate the sediment signature into real catchment processes. This would lead to a better understanding of the different factors contributing to the accumulation of sediment during a year. Once tracing of these factors is achieved a quantitative interpretation of the sediment record will be possible. This calibration, however would only be valid for the interannual variations of MAR in Lej da Silvaplauna. When decadal or centennial scale MAR fluctuations are considered glaciologic parameters exert a more important influence as has been demonstrated in chapter 4. Conclusions 172 Chapter 8

A result that came out clearly from the geochemical and mineralogical analyses presented in chapters 4, 5 and 6 is, that geochemical sediment signatures apparently are not ‘the tracer of choice’ in high alpine proglacial sediments. They are either insensitive to environmental changes or yield results that are equivocal. It was demonstrated that for instance productivity indicators like biogenic silica and organic carbon accumulation do not respond significantly to the severest Holocene cooling event, the LIA. Most likely this is because for this time window the system is totally dominated by detrital glacigenic input that renders any autochthonous signal unde- tectable. Consequently it remains to be tested, if the same is true for the early and middle Holocene climatic optimum when catchment glaciation was much smaller or absent. This then should be done in conjunction with an analysis of biologic parameters. In the early and mid Holocene record of Lej da Silvaplauna e.g. diatoms in varying amounts are present throughout the whole section. Their determination could yield additional information about lake internal changes. Together with changes in the pollen assemblage a much more detailed picture of the early and mid Holocene environmental changes may be drawn then. However, the geochemical analysis in chapter 5 demonstrated that, although differences in sediment bulk chemistry between the early and the late Holocene are detectable, their interpretation remains ambiguous. Moreover, their determination can not realistically be done in a high resolution mode as would be necessary for this study. Hence, geochemical tracers might be useful in sediments that are not dominated by allochthonous input of glacier milk, when combined with other, biological or mineralogical tracers. In purely clastic, almost pristine sediments with more than 98% allochthonous components (chapter 4) they are not to be recommended.

In contrast chapter 5 and 7 showed that the physical properties and the mineralogy of the sediments seem to be well suited as tracers for the catchment glacier history. In chapter 7 changes in the mineralogy of the Late Glacial sediments were interpreted as changes in the positions of the catchment glaciers. It was stressed that the proposed interpretations are not unambiguous yet for two reasons. (1) The climatic interpretation has to be confirmed by independent e.g. palynological evidence. (2) The absolute dating has to be improved for the comparison with other high resolution climate records. These two points are certainly the most important tasks to pursue in the future. Especially, it is very important to establish a detailed pollen stratigraphy, because this, for the moment, is the only source of information about the Late Glacial high alpine climate that is independent from the sedimentological evidence presented in this study. It was mentioned, that such a study is already under way. Chapter 8 173 Conclusions

With this information it will be possible to decide whether the climatic interpretation of the sediment record is correct or has to be modified. Nevertheless, merely the fact that distinct mineralogical changes do occur in the sediments underlines the potential of this sediment signature as tracer for the glacier history in a geologically diverse catchment. It was described in chapter 7 that changes in sediment color occur in the Late Glacial sediments. Most likely these changes reflect the sediment mineralogy as was proposed for the green rhythmites. There- fore a digital grey value or color analysis of the sediment could yield valuable information. These methods have already been applied successfully to laminated marine sediments from the Santa Barbara Basin (e.g. Merrill and Beck, 1995; Schaaf and Thurow, 1994; Schaaf and Thurow, 1995). There sediment color is mostly a function of the total organic carbon (TOC) content and differences in the terrigenous input. Because, TOC content is negligible in the Late Glacial sediments color changes solely should reflect differences in mineralogy. A digital color analysis therefore seems to be the ideal tool to obtain high resolution information about sediment mineralogy. In order to test the proposed method of reconstructing the glacier history with mineralogical tracers the sediments of Lej da Segl would offer an ideal opportunity. A long sediment core is available from the southern basin of this lake (appendix 7.1). Today this basin only receives glacier milk indirectly from the Fedoz catchment through the northern basin of Lej da Segl that is separated by a narrow passage with a sill from the southern basin. But during the Late Glacial the lake also received glacier milk from another catchment (Vadret da Forno) that is lithologically distinctly different. From moraine studies it is known that input from the latter catchment stopped during the Egesen stadial. This certainly left a mineralogical signature in the sediments of this basin that should be detectable very clearly.

Concerning the physical properties more knowledge is needed about the relationship between glacier characteristics, MAR and grainsize median in the proglacial lake sediments. The ‘discrepancy’ of low MAR coinciding with low median grainsize was discussed in chapter 7. But again, the mere occurrence of distinct changes in these two parameters is viewed as evidence for their sensitive reaction to environmental changes. It seems that especially the median grainsize provides a relatively clear signal. Besides being an artefact of the poorer resolution this could be an indication that the median is controlled by one or only a few influencing factors. In contrast it is likely that the much noisier MAR signal is influenced by several factors (chapter 4). An advantage of MAR is that it can be determined with true annual time resolution, provided that the annual nature of lamination was confirmed by 14C- dating. In contrast high resolution grainsize measurements can not realistically be Conclusions 174 Chapter 8 done for 650 cm of sediment because this would involve a minimum of 6000 grainsize measurements. With the technique applied in this study this would take approximately two years of nothing but measuring. One method to obtain high resolution data of the sediment physical properties is microdensitometry of X-ray radiographs, like it is commonly applied in dendro- chronology (Schweingruber, 1989). X-ray densities usually are predominantly controlled by the TOC content of the sediments (Algeo and Woods, 1994; Karlén, 1976; Karlén, 1981). However, given the very low TOC content of the Late Glacial sediments it is believed that X-ray densities rather would reflect variations in wet bulk density (Axelsson, 1983). Wet bulk density in turn is mostly controlled by the water content and thus the porosity of the sediments (chapter 7). How porosity relates to the sediment grainsize remains to be determined (cf. chapter 7). The wet bulk density presented in figure 7.6 was determined by gamma ray attenuation on whole cores (59 mm diameter). This involves an error because this method does not account for tilted or bent sediment layers which is a feature that often occurs in lake sediment cores (cf. figure 7.2). Moreover, the record is discontinuous due to the measuring technique, which could be improved by applying newer techniques. In spite of these drawbacks the wet bulk density record shows variations that are similar to the median grainsize and the water content. Moreover, the high frequency variations also exhibit differences throughout the core. Together this indicates that interannual or interdecadal differences in porosity do exist and thus should be detectable by microdensitometry. A first test series of X-ray radiographs yielded pic- tures where inter- and also intra-annual grey-value changes are clearly visible. How- ever, for an analysis with microdensitometry sediment slabs of precisely constant thickness (preferably 2.0 ± 0.02 mm) are needed. With respect to lake sediments, for the moment this requirement poses a major problem because unlike wood the sediment has to be impregnated with a low viscosity resin prior to cutting with a special saw. But once this problem is solved this method certainly is a powerful tool in high resolution studies of laminated clastic sediments. For instance this method also could make the determination of varves more ‘objective’. Currently varve measurements as conducted in this study (3.4.5) are prone to mis-interpretation by the examiner. This source of error could be reduced if varve boundaries instead of being detected by visual inspection of a biased observer, could be defined by the crossing of a certain X-ray density value. This follows the idea of Leemann and Niessen (1994a) who could define varve boundaries in a 20 cm long sediment section of proglacial Oeschinensee by the decrease of the median below 5 µm in the clay top of each varve. References 175 References

References

Agassiz, L. (1840): Etudes sur les glaciers. Privately published, Neuchâtel. Algeo, T.J. and Woods, A.D. (1994): Microstratigraphy of the Lower Mississippian Sunbury Shale: A record of solar-modulated climatic cyclicity. Geology 22, 795-798. Alley, R.B., Meese, D.A., Shuman, C.A., Gow, A.J., Taylor, K.C., Grootes, P.M., White, J.W.C., Ram, M., Waddington, E.D., Mayewski, P.A. and Zielinski, G.A. (1993): Abrupt increase in Greenland snow accumulation at the end of the Younger Dryas event. Nature 362, 527-529. Ammann, B., Birks, H.J.B., Drescher-Scneider, R., Juggins, S., G., L. and Lotter, A.F. (1993): Patterns of variation in Late-Glacial Pollen stratigraphy along a northwest-southeast transect through Switzerland - A numerical Analysis. Quaternary Science Reviews 12, 277-286. Ammann, B., Lotter, A.F., Eicher, U., Gaillard, M.-J., Wohlfarth, B., Haeberli, W., Lister, G., Maisch, M., Niessen, F. and Schlüchter, C. (1994): The Würmian Late-glacial in lowland Switzerland. Journal of Quaternary Science 9(2), 119-125. Anderson, R.Y. (1993b): The varve chronometer in Elk Lake: Record of climatic variability and evidence for solar-geomagnetic-14C-climate connection. In: Bradbury, J.P. and Dean, W.E. (eds.) Elk Lake, Minnesota: evidence for rapid climate change in the north-central United States, Geological Society of America Special Paper, Vol. 276, pp 45-67, Boulder, Colorado. Anderson, R.Y. and Koopmans, I.H. (1963): Harmonic analysis of varve time series. Journal of Geophysical Research 68, 877-893. Antevs, E. (1930): A Geological Chronometer The Varved Glacial Clays Give an Accurate Measure of the Ages. Canadian Mining Journal 51, 388-390. Ariztegui, D. (1993): Palaeoenvironmental and Palaeoclimatic Implications of Sedimented Organic Matter Variations in Lacustrine Systtems: Lake St. Morirtz, A Case Study. PhD thesis No. 10225, 188p, ETH Zürich. Ariztegui, D., Bianchi, M.M., Masaferro, J., Lafargue, E. and Niessen, F. (1997): Interhemispheric synchrony of Late-glacial climatic instability as recorded in proglacial Lake Mascardi, Argentina. Journal of Quaternary Science 12, 333-338. Ariztegui, D. and Dobson, J. (1996): Magnetic investigations of framboidal greigite formation: a record of anthropogenic environmental changes in eutrophic Lake St. Mortiz, Switzerland. The Holocene 6(2), 235-241. Ariztegui, D., Niessen, F. and von Salis, K. (1992): Lesen wie im Geschichtsbuch: Seenforschung im St. Moritzer See. Engadiner Post 7, St. Moritz. Arrhenius, S. (1896): On the influence of Carbonic Acid in the Air upon the Temperature of the Ground. Philosophical Magazine 41(237). Attolini, M.R., Galli, M. and Nanni, T. (1988): Long and short cycles in solar activity during the last millennia. In: Stephenson, F.R. and Wolfendale, A.W. (eds.) Secular Solar and Geomagnetic Variations in the Last 10000 Years, Vol. 236, pp 49-68, Kluwer Academic Publishers, Dordrecht, etc.. Axelsson, V. (1983): The use of X-ray radiographic methods in studying sedimentary properties and rates of sediment accumulation. Hydrobiologia 103, 65-69. Becker, B., Kromer, B. and Trimborn, P. (1991): A stable-isotope tree ring timescale of the Late Glacial/ Holocene boundary. Nature 353, 647-649. Behre, K.-E. (1978): Die Klimaschwankungen im europäischen Präboreal. Petermanns Geographische Mitteilungen 122(2), 97-102. References 176 References

Beniston, M., Diaz, H.F. and Bradley, R.S. (1997): Climatic change at high elevation sites: An overview. Climatic Change 36(3-4), 233-251. Beniston, M., Rebetez, M., Giorgi, F. and Marinucci, M.R. (1994): An Analysis of Regional Climate Change in Switzerland. Theoretical and Applied Climatology 49, 135-159. Berner, R.A. (1971): Principles of Chemical Sedimentology. McGraw-Hill Book Company, 240p, New York. Berner, R.A. (1990): Atmospheric carbon dioxide levels over Phanerozoic time. Science 249, 1382-1386. Berner, R.A. and Lasaga, A.C. (1989): Simulation des geochemischen Kohlenstoffkreislaufs. Spektrum der Wissenschaft Mai ‘89, 54-62. Bingham Müller, T.L. (1996): Microbially Mediated Iron Mineral Transformations: A Case Study in Lake Greifen, Switzerland. Ph.DThesis No. 11752, ETH Zürich. Biondi, F., Lange, C.B., Huges, M.K. and Berger, W.H. (1997): Inter-decadal signals during the last millenium (AD 1117-1992) in the varve record of Santa Barbara basin, California. Geophysical Research Letters 24(2), 193-196. Björck, S., Kromer, B., Johnsen, S., Bennike, O., Hammarlund, D., Lemdahl, G., Possnert, G., Rasmussen, T.L., Wohlfarth, B., Hammer, C.U. and Spurk, M. (1996): Synchronized Terrestrial-Atmospheric Deglacial Records Around the North Atlantic. Science 274, 1155-1160. Borner, L. (1922): Die Bodenfauna des St. Moritzer Sees. Eine monographische Studie. Archiv für Hydrobiologie 13. Bosli-Pavoni. (1971): Ergebnisse der limnologischen Untersuchungen der Oberengadiner Seen. Schweizerische Zeitschrift zur Hydrologie 33, 386-409. Bouma, A.H. (1962): Sedimentology of some Flysch Deposits. A graphic approach to facies interpretation. Elsevier Publishing company, 168p, Amsterdam, New York. Brown, G.H., Tranter, M. and Sharp, M.J. (1996): Experimental investigations of the weathering of suspended sediment by alpine glacial meltwater. Hydrological Processes 10, 579-597. Brückner, E. (1889): In wie weit ist das heutige Klima konstant ? Verhandlungen des VIII. Deutschen Geographentages, pp 101-115, Verlag von Dietrich Reimer, Berlin. Burga, C.A. (1988): Swiss vegetation history during the last 18000 years. New Phytologist 110, 581-602. Burga, C.A., Perret, R., Gehrig, S., Vogel, H.-H., Maggetti, B., Fitze, P., Maisch, M. and Graf, K. (1997): Geoökologische und klimamorphologische Untersuchungen im Alpenraum. Isola-Delta, Val Fedoz und Val Fex, Oberengadin, Kanton Graubünden. vdf Hochschulverlag AG an der ETH Zürich, 57p, Zürich. BUS. (1985): Gewässerschutzstatistik. Schriftenreihe Umweltschutz 46, Bundesamt für Umweltschutz, Bern. Çambel, A.B. (1993): Applied Caos Theory. A paradigm for complexity. Academic Press, Inc., 246p, Boston, etc.. Collins, D.N. (1990): Seasonal and annual variations of suspended sediment transport in meltwaters draining from an Alpine glacier. IAHS Publications 193, 439-446. Collins, D.N. and Taylor, D.P. (1990): Variability of Runoff from Partially-Glacierized Alpine Basins. IAHS Publications 193, 365-373. Conrad, H. (1981): Schriften zur urgeschichtlichen und römischen Besiedlung des Engadins. Engadin Press AG, Samedan, 138p. Cornelius, H.P. (1932): Geologische Karte der Err-Julier-Gruppe, Ost- und Westblatt, Spezialkarte 115 (1:50000). Schweizerische Geologische Kommission. Covey, C. and Hoffert, M.I. (1997): The Sun-Climate Connection: A Challenge to Conventional Wisdom; An Editorial Essay. Climatic Change 37(2), 387-390. References 177 References

Creer, K.M. and Tucholka, P. (1983): On the current state of lake sediment palaeomagnetic research. Geophys. J. R. astr. Soc. 74, 223-238. Crowley, T.J. and North, G.R. (1991): Paleoclimatology. Oxford University Press, 339p, New York. Dansgaard, W., Johnsen, S.J., Clausen, H.B., Dahl-Jensen, D., Gundestrup, N.S., Hammer, C.U., Hvidberg, C.S., Steffensen, J.P., Sveinbjörnsdottir, A.E., Jouzel, J. and Bond, G. (1993): Evidence for general instability of past climate from a 250-year ice-core record. Nature 364, 218-220. Dansgaard, W., White, J.W.C. and Johnsen, S.J. (1989): The abrupt termination of the Younger Dryas climate event. Nature 339, 532-534. de Charpentier, M.J. (1835): Notice sur la Cause probable du Transport des Blocs Erratiques de la Suisse. Carilian-Goeury, Libraire-Editeur, pp 1-20, Paris. de Charpentier, M.J. (1841): Essai sur les glaciers et le terrain erratique du bassin du Rhone. Marc Ducloux, pp 241-248, Lausanne. de Geer, G. (1912): A geochronology of the last 12,000 years. XI Internat. Geol. Congr. Stockholm 1910 compte rendue 1, 241-258. Dean, J., W.E. and Anderson, R.Y. (1974): Application of Some Correlation Coefficient Techniques to Time-Series Analysis. Mathematical Geology 6(4), 363-372. Dean, W.E. (1993): Physikal properties, mineralogy and geochemistry of Holocene varved sediments from Elk Lake, Minnesota. In: Bradbury, J.P.D., W.E. (eds.) Elk Lake, Minnesota: evidence for rapid climate change in the north-central United States, Geological Society of America Special Paper, Vol. 276, pp135-157, Boulder, Colorado. Dearing, J.A. (1991): Lake sediment records of erosional processes. Hydrobiologia 214, 99-106. Dearing, J.A. and Flower, R.J. (1982): The magnetic susceptibility of sedimenting material trapped in Lough Neagh, Northern Ireland, and its erosional significance. Limnology and Oceanography 27(5), 969-975. Dearing, J.A. and Foster, I.D.L. (1986): Limnic sediments used to reconstruct sediment yields and sources in the English Midlands since 1765. In: Gardiner, V. (ed.) International Geomorphology I, pp 853-868, John Wiley & Sons. Delworth, T., Manabe, S. and Stouffer, R.J. (1993): Interdecadal Variations of the Thermohaline Circulation in a Coupled Ocean-Atmosphere Model. Journal of Climate 6(11), 1993-2011. Desloges, J.R. (1994): Varve Deposition and the Sediment Yield Record at Three Small Lakes of the Southern Canadian Cordillera. Arctic and Alpine Research 26(2), 130-140. Dhillon, K.S., Sinha, M.K. and Randhawa, N.S. (1975): Organo-Metallic Phosphates V. Complexation of Zinc and Phosphorus by Humic Compounds. Plant and Soil 43, 317-326. Dietrich, D. (1971): Die Ophiolithe am Piz Corvatsch (Oberengadin). Diploma-Thesis, 80p, ETH Zürich. Dietrich, V. (1970): Die Stratigraphie der Platta-Decke. Fazielle Zusammenhänge zwischen Oberpenninikum und Unterostalpin. Eclogae Geologicae Helvetiae 63(2), 631-671. Drever, J.I. and Hurcomb, D.R. (1986): Neutralization of atmospheric acidity by chemical weathering in an alpine drainage basin in the North Cascade Mountains. Geology 14, 221-224. Drever, J.I. and Zobrist, J. (1992): Chemical weathering of silicate rocks as a function of elevation in the southern Swiss Alps. Geochimica et Cosmochimica Acta 56, 3209-3216. Eicher, U. and Siegenthaler, U. (1976): Palynological and oxygen isotope investigations on Late-Glacial sediment cores from Swiss lakes. Boreas 5, 109-117. Engstrom, D.R. and Wright Jr., H.E. (1984): Chemical stratigraphy of lake sediments as a record of environmental change. In: Haworth, E.Y. and Lund, J.W.G. (eds.) Lake Sediments and Environmental History, pp 11-67, Leicester University Press, Leicester. Epstein, S. (1995): The isotopic climatic records in the Alleröd-Bølling-Younger Dryas and post-Younger Dryas events. Global Biogeochemical Cycles 9(4), 557-563. References 178 References

Escher von der Linth, A. and Studer, B. (1839): Geologische Beschreibung von Mittel-Bündten. Denkschriften der Schweizerischen Naturforschenden Gesellschaft 3(2), 218. Fanning, D.S., Keramidas, V.Z. and El-Desoky, M. (1989): Micas. In: Dixon, J.B. and Weed, S.B. (eds.) Minerals in Soil Environments, pp 551-634, Soil Science Society of America, Madison, Wisconsin. Fenn, C.R. and Gomez, B. (1989): Particle size analysis of the sediment suspended in a proglacial stream: Glacier de Tsidjiore Nouve, Switzerland. Hydrological Processes 3, 123-135. Finger, W., Mercolli, I., Kündig, R., Stäubli, A., de Capitani, C., Nievergelt, P., Peters, T. and Trommsdorff, V. (1982): Bericht über die gemeinsame Exkursion der Schweizerischen Geologischen Gesellschaft und der Schweizerischen Mineralogischen und Petrographischen Gesellschaft ins Oberengadin vom 21. bis 24. September 1981. Eclogae geologicae Helvetiae 75(1), 199-222. Fischer, A. (1996): Isotopengeochemische Untersuchungen (∂18O und ∂13C) im Wasser und in den Sedimenten des Soppensees (Kt. Luzern, Schweiz). Klimaveränderungen und Entwicklungsgeschichte des Sees seit dem Spätglazial. Ph.D. Thesis, No. 11924, 184p, ETH Zürich. Fliri, F., Bortenschlager, S., Felber, H., Heissel, W., Hilscher, H. and Resch, W. (1970): Der Bänderton von Baumkirchen (Inntal, Tirol). Eine neue Schlüsselstelle zur Kenntnis der Würm-Vereisung der Alpen. Zeitschrift für Gletscherkunde und Glazialgeologie 6(1-2), 5-35. Föllmi, K.B. (1995): 160 m.y. record of marine sedimentary phosphorus burial: Coupling of climate and continental weathering under greenhouse and icehouse conditions. Geology 23(9), 859-862. Förstner, U. (1992): Umweltschutztechnik. Springer-Verlag, 507p, Berlin. Förstner, U. and Wittmann, G.T.W. (1979): Metal Pollution in the Aquatic Environment. Springer Verlag, pp 197-238, Berlin, Heidelberg, New York. Fountain, A.G. and Tangborn, W.V. (1985): The effect of glaciers on streamflow variations. Water Resources Research 21, 579-586. Furrer, G., Burga, C., Gamper, M., Holzhauser, H.-P. and Maisch, M. (1987): Zur Gletscher-, Vegetations- und Kimageschichte der Schweiz seit der Späteiszeit. Geographica Helvetica 2, 61-91. Furrer, G., Gamper-Schollenberger, B. and Suter, J. (1980): Zur Geschichte unserer Gletscher in der Nacheszeit - Methoden und Ergebnisse. In: Oeschger, H., Messerli, B. and Svilar, M. (eds.) Das Klima. Analysen und Modelle Geschichte und Zukunft, pp 91-107, Springer-Verlag, Berlin Heidelberg New York. Gamper, M. (1982): Postglaziale Solifluktionsphasen am Albulapass (östliche Schweizer Alpen). Physische Geographie 1, 171-186. Gamper, M. and Suter, J. (1978): Der Einfluss von Temperaturänderungen auf die Länge von Gletscherzungen. Geographica Helvetica 4, 183-189. Gehrig, S. (1996): Stratigraphische, pedologische und geomorphodynamische Untersuchungen am Isola-Delta. Abfolge, Intensität und Häufigkeit von Ereignissen. Diplomarbeit, Geographisches Institut der Universität Zürich. Gensler, G.A. (1978): Das Klima von Graubünden. Ein Beitrag zur Regionalklimatologie der Schweiz. Habilitationsschrift, Universität Zürich. Gerland, S. (1993): Zerstörungsfreie hochauflösende Dichteuntersuchungen mariner Sedimente. Berichte zur Polarforschung 123, 1-129. Ghil, M. and Vautard, R. (1991): Interdecadal oscillations and the warming trend in global temperature time series. Nature 350, 324-327. Gilbert, R. (1975): Sedimentation in Lilooet Lake, British Columbia. Canadian Journal of Earth Science 12(10), 1697-1711. Gilly, M.R. (1992): Wasserversorgung Maloja. Grundwasseruntersuchungen bei Isola. Gutachten im Auftrag der Gemeinde Stampa, unpublished. Giovanoli, F. (1979): Die remanente Magnetisierung von Seesedimenten. Mitt. Geol. Inst. ETH Zürich 230, pp 1-200. References 179 References

Giovanoli, F., Kelts, K.R., Finckh, P. and Hsü, K.J. (1984): Geological framework, site survey and seismic stratigraphy. In: Hsü, K.J.K., K.R. (eds.) Quaternary Geology of Lake Zürich: An Interdisciplinary Investigation by Deep-Lake Drilling, Contributions to Sedimentology 13, pp 5-20. Giovanoli, R., Schnoor, J.L., Sigg, L., Stumm, W. and Zobrist, J. (1988): Chemical weathering of crystalline rocks in the catchment area of acidic Ticino Lakes, Switzerland. Clays and Clay Minerals 36(6), 521-529. Goethe, J.W. (1829): Wilhelm Meisters Wanderjahre 2tes Buch Cap. 10, Goethes Werke Bd. 22, Cotta. Goslar, T., Arnold, M., Bard, E., Kuc, T., Pazdur, M.F., Ralska-Jasiewiczowa, M., Rozanski, K., Tisnerat, N., Walanus, A., Wicik, B. and Wieckowski, K. (1995): High concentration of atmospheric 14C during the Younger Dryas cold episode. Nature 377, 414-417. Goslar, T., Kuc, T., Ralska-Jasiewiczowa, M., Rozanski, K., Arnold, M., Bard, E., van Geel, B., Pazdur, M.F., Szeroczynska, K., Wicik, B., Wieckowski, K. and Walanus, A. (1993): High-resolution lacustrine record of the Late Glacial/Holocene transition in central Europe. Quaternary Science Reviews 12, 287-294. Granar, L. (1956): Dating of recent fluvial sediments from the estuary of the Ångerman river. Geologiska Föreningens i Stockholms Förhandlingar 78(4), 654-658. Gross, E. and Ritzmann, C. (1990): Die neolithischen und bronzezeitlichen Siedlungen im Züricher Seefeld. Die ersten Bauern, Pfahlbaufunde Europas, Band 1: Schweiz, pp 161-171, Zürich. Gross, G., Kerschner, H. and Patzelt, G. (1976): Methodische Untersuchungen über die Schneegrenze in alpinen Gletschergbieten. Zeitschrift für Gletscherkunde und Glazialgeologie 12(2), 223-251. Habbe, K.A. (1996): überlegungen zum Bewegungsmechanismus vorstossender kaltzeitlicher Gletscher und zur glazialen Erosion und Übertiefung. Eclogae geol. Helv. 89(3), 1007-1022. Haeberli, W. (1991): Zur Glaziologie der letzteiszeitlichen Alpenvergletscherung. In: Frenzel, B. (ed.) Klimageschichtliche Probleme der letzten 130000 Jahre 1, pp 409-419. Haeberli, W. (1994): Grundzüge der Glaziologie - Gletscher und Permafrost. Vorlesungsskript ETHZ WS 1993/94, 32p. Haeberli, W. (1995): Glaziale und periglaziale Morphodynamik. Vorlesungsskript ETHZ SS 1995, 32p. Haeberli, W., Guodong, C., Gurbunov, A.P. and Harris, S.A. (1993): Mountain Permafrost and Climatic Change. Permafrost and Periglacial Processes 4, 165-174. Hajdas, I., Ivy, S.D., Beer, J., Bonani, G., Imboden, D., Lotter, A.F., Sturm, M. and Suter, M. (1993): AMS radiocarbon dating and varve chronology of Lake Soppensee: 6000 to 12000 14C years BP. Climate Dynamics 9, 107-116. Hantke, R. (1983): Eiszeitalter; Die jüngste Erdgeschichte der Schweiz und ihrer Nachbargebiete. Ott Verlag Thun, pp 79-196, Switzerland. Hasholt, B.W., D. E. (1992): Use of caesium-137 to investigate sediment sources and sediment delivery in a small glacierized mountain drainage basin in eastern Greenland. IAHS Publications 209, 87- 100. Heim, A. (1879a): Die Seen des Oberengadin. Jahrbuch des Schweizer Alpen Club 15, 429-436. Heinrichs, H. (1990): Aufschlußverfahren in der analytischen Geochemie (Teil 1). Lab. Prax. 12/89, 1-6. Heinrichs, H., Brumsack, H.-J., Loftfield, N. and König, N. (1986): Verbessertes Druckaufschlussystem für biologische und anorganische Materialien. Z. Pflanzenernähr. Bodenk. 149, 350-353. Heinrichs, H. and Herrmann, A.G. (1990): Praktikum der Analytischen Geochemie. Springer-Verlag, 669p, Berlin, Heidelberg, New York. Heitz, A., Punchakunnel, P. and Zoller, H. (1982a): Vegetations, klima und Gletschergeschichte des Oberengadins. In: Gamper, M. (eds.) Beiträge zur Quartärforschung in der Schweiz, Vol. 1, pp 157- 170, Geographisches Institut Universität Zürich, Zurich. References 180 References

Heitz, A., Punchakunnel, P. and Zoller, H. (1982b): Zum Problem der 14C-Datierung im Veltlin und Oberengadin. In: Gamper, M. (eds.) Beiträge zur Quartärforschung in der Schweiz, Vol. 1, pp 91-102, Geographisches Institut Universität Zürich, Zurich. Heller, F., Xiuming, L., Tungsheng, L. and Tongchun, X. (1991): Magnetic susceptibility of loess in China. Earth and Planetary Science Letters 103, 301-310. Heuberger, H. (1968): Die Alpengletscher im Spät- und Postglazial. Eine chronologische Übersicht. Eiszeitalter und Gegenwart 19, 270-276. Higgitt, S.R., Oldfield, F. and Appleby, P.G. (1991): The record of land use change and soil erosion in the late Holocene sediments of the Petit Lac d’Annecy, eastern France. The Holocene 1(1), 14-28. Hilgers, E., Thies, H. and Kalbfus, W. (1993): A lead-210 dated sediment record on heavy metals, polycyclic aromatic hydrocarbons and soot spherules for a dystrophic mountain lake. Verh. Internat. Verein. Limnol. 25, 1091-1094. Hoelzle, M. (1994): Permafrost und Gletscher im Oberengadin, Grundlagen und Anwendungsbeispiele für automatisierte Schätzverfahren. Mitteilungen der Versuchsanstalt für Wasserbau, Hydrologie und Glaziologie der ETHZ 132, 121p. Hoelzle, M. and Haeberli, W. (1995): Simulationg the effects of mean annual air-temperature changes on permafrist distribution and glacier size: an example from the Upper Engadin, Swiss Alps. Annals of Glaciology 21, 399-405. Holtmeier, F.-K. (1966): Die “Malojaschlange” und die Verbreitung der Fichte. Beobachtungen zur Klimaökologie des Oberengadins. Wetter und Leben 18(5/6), 105-108. Holtmeier, F.-K. (1967a): Die Waldgrenze im Oberengadin in ihrer physiognomischen und okologischen differenzierung. Rheinische Friedrich-Wilhelms-Universitat., 164p, Bonn. Holtmeier, F.-K. (1967b): Die Verbreitung der Holzarten im Oberengadin unter dem Einfluss des Menschen und des Lokalklimas. Erdkunde 21(4), 249-258. Houghton, J.T., Jenkins, G.T. and Ephraums, J.J. (1990): Climate Change: the IPCC scientific assessment. Cambridge University Press, Cambridge, etc.. Hutchinson, G.E. and Löffler, H. (1956): The thermal classification of lakes. Proceedings of Nat. Acad. Sci. Wash. 42, 84-86. IPCC. (1995b): The Economic and Social Dimensions of Climate Change - IPCC Working Group III. Cambridge Univ. Press, Cambridge. Itkonen, A. and Salonen, V.-P. (1994): The response of sedimentation in three varved lacustrine sequences to air temperature, precipitation and human impact. Journal of Paleolimnology 11, 323- 332. Ivy-Ochs, S., Schlüchter, C., Kubik, P.W., Synal, H.-A., Beer, J. and Kerschner, H. (1996): The exposure age of an Egesen moraine at Julier Pass, Switzerland, measured with the cosmogenic radionuclides 10Be, 26Al and 36Cl. Eclogae geologicae Helveticae 89(3), 1049-1063. Jäckli, H. (1962): Die Vergletscherung der Schweiz im Würmmaximum (mit einer Karte, Tafel 1). Eclogae Geologicae Helvetiae 55(2), 285-294. Jerikowic, J.L. and Sonett, C.P., &, Stihler, S.D., Stone, D.B. and Beget, J.E. (1993): “Varve” counting vs. tephrochronology and 137Cs and 210Pb dating: A comparative test at Skilak Lake, Alaska: Comment and Reply. Geol 21, 763-764. Johnson, S.J., Clausen, H.B., Dansgaard, W., Fuhrer, K., Gundestrup, N., Hammer, C.U., Iversen, P., Jouzel, J., Stauffer, B. and Steffensen, J.P. (1992): Irregular glacial interstadials recorded in a new Greenland ice core. Nature 359, 311-313. Jones, J.G. (1982): Activities of Aerobic and Anaerobic Bacteria in Lake Sediments and their Effect on the Water Column. In: Nedwell, D.B. and Brown, C.M. (eds.) Sediment Microbiology, pp 107-145, Academic Press, London. References 181 References

Jones, P.D. and Bradley, R.S. (1995): Climatic variations in the longest instrumental records. In: Bradley, R.S. and Jones, P.D. (eds.) Climate since A.D. 1500, pp 246-268, Routledge, London, New York. Jones, P.D. and Wigley, T.M.L. (1990): Die Erwärmung der Erde seit 1850. Spektrum der Wissenschaft Okt. 90, 108-116. Karlén, W. (1976): Lacustrine sediments and tree-limit variations as indicators of Holocene climatic fluctuations in Lappland, Northern Sweden. Geografiska Annaler 58 A(1-2), 1-34. Karlén, W. (1981): Lacustrine sediment studies. A technique to obtain a continuous record of Holocene glacier variations. Geografiska Annaler 63 A(3-4), 273-281. Kelts, K., Briegel, U., Ghilardi, K. and Hsü, K. (1986): The limnogeology-ETH coring system. Schweiz. Z. Hydrol. 48(1), 104-115. King, T. (1996): Quantifying nonlinearity and geometry in time series of climate. Quaternary Science Reviews 15, 247-266. Kleiber, H. (1974): Pollenanalytische Untersuchungen zum Eisrückzug und zur Vegetationsgeschichte im Oberengadin I. Botanische Jahrbucher 94(1), 1-53. Kromer, B. and Becker, B. (1993): German oak and pine 14C calibration, 7200-9439 BC. Radiocarbon 35, 125-135. Lang, C. (1885): Der säculare Verlauf der Witterung als Ursache der Gletscherschwankungen in den Alpen. Zeitschrift der Österreichischen Gesellschaft für Meteorologie 20, 443-457. Lauriol, B.M., Lalonde, A.E. and Dewez, V. (1997): Weathering of Quartzite on a Cryoplanation Terrace in Northern Yukon, Canada. Permafrost and Periglacial Processes 8, 147-153. Leemann, A. (1993): Rhythmite in Alpinen Vorgletscherseen: Warvenstratigraphie und Aufzeichnung von Klimaveränderungen. Ph.D. thesis No. 10386, 129p, ETH Zürich. Leemann, A. and Niessen, F. (1994a): Varve formation and the climatic record in an Alpine proglacial lake. The Holocene 4(1), 1-8. Leemann, A. and Niessen, F. (1994b): Holocene glacial activity and climatic variations in the swiss Alps: reconstructing a continuous record from proglacial lake sediments. The Holocene 4(3), 259- 268. Lehman, S.J. and Keigwin, L.D. (1992): Sudden changes in North Atlantic circulation during the last deglaciation. Nature 356, 757-762. Leonard, E.M. (1985a): Glaciological and climatic controls on lake sedimentation, Canadian Rocky Mountains. Z. Gletscherk. Glazialgeol. 21, 35-42. Leonard, E.M. (1985b): Use of Lacustrine Sedimentary Sequences as Indicators of Holocene Glacial History, Banff National Park, Alberta, Canada. Quaternary Research 26, 218-231. Leonard, E.M. (1986): Varve studies in Hector Lake, Alberta, Canada, and the relationship between glacial activity and sedimentation. Quat. Res. 25, 199-214. LIMNEX. (1994): Gewässerzustand und Gewässerschutzmassnahmen im Oberengadin. Bericht zuhanden des Amtes für Umweltschutz, Kanton Graubünden, 75p. Liniger, M. and Guntli, P. (1988): Bau und Geschichte des zentralen Teils der Margna Decke. Schweizerische Mineralogische und Petrographische Mitteilungen 68, 41-54. Lins, H.F. (1994): Recent Directions Taken in Water, Energy, and Biogeochemical Budgets Research. Transactions, American Geophysical Union 75(38), 433-439. Lister, G.S. (1984): Deglaciation of the Lake Zürich area: a model based on the sedimentological record. Contributions to sedimentology 13, 177-185. Lister, G.S. (1985): Late Pleistocene Alpine Deglaciation and Post-Glacial Climatic Developements in Switzerland: The Record from Sediments in a Peri-Alpine Lake Basin. Ph.D. Thesis No. 7753, 151p, ETH Zürich. References 182 References

Lister, G.S. (1988): A 15000-Year Isotopic record from Lake Zürich of Deglaciation and Climatic Change in Switzerland. Quaternary Research 29, 129-141. Livingstone, D.M. (1997): Break-Up Dates of Alpine Lakes as Proxy Data for Local and Regional Mean Surface Air Temperatures. Climatic Change 37(2), 407-439. Livingstone, D.M. and Hajdas, I. (1998 in press): Climatically relevant periodiceties in the thickness of biogenic carbonate varves in Soppensee, 9740 - 6870 calendar yr BP. Journal of Paleolimnology . Lotter, A.F., Ammann, B., Beer, J., Hajdas, I. and Sturm, M. (1992): A step towards an absolute timescale for the Late-Glacial: Annually laminated sediments from Soppensee (Switzerland). In: Bard, E. and Broecker, W.S. (eds.) The last deglaciation: Absolute and Radiocarbon Chronologies, Vol. 12, pp 45-68, Springer-Verlag, Berlin, Heidelberg. Lotter, A.F., Ammann, B. and Sturm, M. (1992): Rates of change and chronological problems during the late-glacial period. Climate Dynamics 6, 233-239. Lotter, A.F., Eicher, U., Siegenthaler, U. and Birks, H.J.B. (1992): Late-Glacial climatic oscillations as recorded in Swiss lake sediments. Journal of Quaternary Science 7, 187-204. Lütschg-Loetscher, O. (1954): Die Eis- und Schneeverhältnisse der Oberengadiner Seen insbesondere des St. Moritzer Sees. Beitrag zur Gewässer- und Kimakunde des Oberengadins. Zum Wasserhaushalt des Schweizer Hochgebirges 1. Bd., III Teil, 10. Kapitel, pp 1-162. Maggetti, B. (1994): Geomorphologische und Geochronologische sowie Gletschergeschchtliche Untersuchungen im Val Fex (Graubünden). Diploma thesis, 108p, Geographical Institute University of Zürich. Mair, M. (1995): Die Oberengadiner Seen. Verlag Paul Haupt, 95p, Bern. Maisch, M. (1981): Glazialmorphologische und Gletschergeschichtliche Untersuchungen im Gebiet zwischen Landwasser- und Albulatal (Kt. Graubünden, Schweiz). Physische Geographie 3, 215p. Maisch, M. (1992): Die Gletscher Graubündens. Habil Schrift Geogr. Inst. Univ. Zürich, Teil A und B, 428p Physische Geographie 33. Maisch, M. and Haeberli, W. (1982): Interpretation geometrischer Parameter von Spätglazial-Gletschern im Gebiet Mittelbünden, schweizer Alpen. In: Gamper, M. (ed.) Beiträge zur Quartärforschung in der Schweiz, Vol. 1, pp 111-126, Geographisches Institut Universität Zürich, Zurich. Makereth, F.J.H. (1966): Some chemical observations on post-glacial lake sediments. Phil. Trans. R. Soc. B. 250, 165-213. Mangerud, J., Andersen, S.T., Berglund, B.E. and Donner, J.J. (1974): Quaternary stratigraphy of Norden, a proposal for terminology and classification. Boreas 3, 109-128. Manley, G. (1974): Central England temperatures: Monthly means 1659-1972. Quarterly Journal of the Royal Meteorological Society of London 100, 389-405. Mann, M.E., Bradley, R.S. and Huges, M.K. (1998): Global scale temperature patterns and climate forcing over the past six centuries. Nature 392, 779-787. Märki, E. (1966): Vorbericht über die chemisch-physikalische Untersuchung der Oberengadiner Seen vom 30. August - 1. September 1965. Bericht zuhanden der internationalen Arbeitsgemeinschaft für die Donauforschung . Mast, M.A. and Drever, J.I. (1990): Chemical Weathering in the Loch Vale Watershed, Rocky Mountain National Park. Colorado. Water Resources Research 26(12), 2971-2978. McManus, J. (1988): Grain size determination and interpretation. In: Tucker, M. (ed.) Techniques in Sedimentology, pp 63-85, Blackwell Scientific Publications, Oxford, etc. Menounos, B. (1997): The water content of lake sediments and its relationship to other physical parameters: an alpine case study. The Holocene 7(2), 207-212. Merrill, R.B. and Beck, J.W. (1995): The ODP color digital imaging system: Color logs of Quaternary sediments from the Santa Barbara Basin, site 893. In: Kennett, J.P., Baldauf, J.G. and Lyle, M. (eds.) Proceedings of the Ocean Drilling Program, Scientific Results (Part 2), Vol. 146, pp 45-59, Ocean Drilling Program, Texas A&M University. References 183 References

Messerli, B., Messerli, P., Pfister, C. and Zumbühl, H.J. (1978): Fluctuations of climate and glaciers in the Bernese Oberland, Switzerland and their geoecological significance, 1600-1975. Arctic and Alpine Research 10, 247-260. Müller, P.J. and Schneider, R. (1993): An automated leaching method for the determination of opal in sediments and particulate matter. Deep Sea Research 40(3), 435-444. Naef, H. (1987): Ein Beitrag zur Stratigraphie der Trias-Serien im Unterostalpin Graubündens (Grisoniden). Mitteilungen aus dem Geologischen Institut der Eidgenössischen Technischen Hochschule und der Universität Zürich 276, 206p. Nesje, A. and Dahl, S.O. (1993): Late Glacial and Holocene Glacier Fluctuations and Climate Variations in Western Norway: A Review. Quaternary Science Reviews 12(4), 255-261. Nesje, A., Johannessen, T. and Birks, H.J.B. (1995): Briksdalsbreen, western Noorway: climatic effects on terminal response of a temperate glacier between 1901 and 1994. The Holocene 5(3), 343-347. Niessen, F. (1996): Physical Properties in Marine Sediments. In: Kuhn, G. (ed.) Die Expedition ANTARKTIS-XI/4 MIT FS”POLARSTERN” 1994, Berichte zur Polarforschung. Niessen, F. and Kelts, K. (1989): The deglaciation and Holocene sedimentary evolution of southern perialpine Lake Lugano-implications for Alpine paleoclimate. Eclogae Geologicae Helvetiae 82(1), 235-263. Niessen, F., Lami, A. and Guilizzoni, P. (1993): Climatic and tectonic effects on sedimentation in central Italian volcano lakes (Latium) - Implications from high resolution seismic profiles. In: Negendank, J.F.W. and Zolitschka, B. (eds.) Paleolimnology of European Maar Lakes, Vol. 49, pp 129-148, Springer Verlag, Berlin, Heidelberg, New York. Niessen, F., Wick, L., Bonani, G., Chondrogianni, C. and Siegenthaler, C. (1992): Aquatic system response to climatic and human changes: productivity, bottom water oxygen status, and sapropel formation in Lake Lugano over the last 10 000 years. Aquatic sciences 54(3/4), 257-276. Nowaczyk, N.R. (1991): Hochauflösende Magnetostratigraphie spätquartärer Sedimente arktischer Meeresgebiete. Berichte zur Polarforschung 78, 1-187. Ohlendorf, C. (1993): Untersuchungen zur Eutrophierungs- und Belastungsgeschichte und Kartierung der Sedimente des Grossen Montiggler Sees (Nord-Italien). Diplomarbeit, 127p, Georg-August- Universität, Göttingen. Ohlendorf, C., Leemann, A. and Niessen, F. (1995): High alpine lake sediments as continuous archives recording glacier and climate history. INQUA XIV International Congress, Schriften der Alfred-Wegener- Stiftung 2, 203. Ohlendorf, C., Niessen, F. and H., W. (1997): Glacial Varve Thickness and 127 Years of Instrumental Climate Data: A Comparison. Climatic Change 36(3-4), 391-411. Oldfield, F. and Clark, R.L. (1990): Lake Sediment-based Studies of Soil Erosion. In: Boardman, J., Foster, I.D.L. and Dearing, J.A. (eds.) Soil Erosion on Agricultural Land, pp 201-228, John Wiley & Sons, Chichester, NewYork, etc. Oldfield, F., Hunt, A., Jones, M.D.H., Chester, R., Dearing, J.A., Olsson, L. and Prospero, J.M. (1985a): Magnetic differentiation of atmospheric dusts. Nature 317, 516-518. Oldfield, F., Maher, B.A., Donoghue, J. and Pierce, J. (1985b): Particle-size related, mineral magnetic source sediment linkages in the Rhode River catchment, Maryland, USA. Journal of the Geological Society of London 142, 1035-1046. Olsson, I.U. (1979): A warning against radiocarbon dating of samples containing little carbon. Boreas 8(1), 203-207. Østrem, G. (1975): Sediment transport in glacial meltwater streams. In: Jopling, A.V. and McDonald, B.C. (eds.) Glaciofluvial and glaciolacustrine sedimentation, pp 101-122. Paillard, D., Labeyrie, L. and Yiou, P. (1996): Macintosh program performs time-series analysis. EOS Trans. AGU 77, 379. References 184 References

Parker, D.E., Legg, T.P. and Folland, C.K. (1992): A New Daily Central England Temperature Series, 1772-1991. Int. J. Climatol. 12, 317-342. Paterson, W.S.B. (1994): The Physics of Glaciers. Elsevier Science Ltd, 480p, Oxford. Patzelt, G. (1972): Die spätglazialen Stadien und postglazialen Schwankungen von Ostalpengletschern. Berichte der Deutschen Botanischen Gesellschaft 85(1-4), 47-57. Patzelt, G. (1985): The Period of Glacier Advances in the Alps 1965 to 1980. Zeitschrift für Gletscherkunde und Glazialgeologie 21, 403-407. Patzelt, G. (1977): Der zeitliche Ablauf und das Ausmass postglazialer Klimaschwankungen in den Alpen. In: Frenzel, B. (eds.) Dendrochronologie und Postglaziale Klimaschwankungen in Europa, pp 248-259, Franz Steiner Verlag GmbH, Wiesbaden. Pavoni, N. (1977): Erdbeben im Gebiet der Schweiz. Eclogae Geologicae Helvetiae 70(2), 351. Penck, A. and Brückner, E. (1909): Die Alpen im Eiszeitalter. Chr. Herm. Tauchnitz, 393p, Leipzig. Perkins, J.A. and Sims, J.D. (1983): Correlation of Alaskan varve thickness with climatic parameters and use in palaeoclimatic reconstruction. Quat. Res. 20, 308-321. Peteet, D. (1995): Global Younger Dryas ? Quaternary International 28, 93-104. Pfister, C. (1980): Klimaschwankungen und Witterungsverhältnisse im schweizerischen Mittelland und Alpenvorland zur Zeit des “Little Ice Age”. Die Aussage der historischen Quellen. In: Oeschger, H., Messerli, B. and Svilar, M. (eds.) Das Klima. Analysen und Modelle Geschichte und Zukunft, pp 175-190, Springer-Verlag, Berlin, etc. Pfister, C. (1995): Monthly temperature and precipitation in central Europe 1525-1979: quantifying documentary evidence on weather and its effects. In: Bradley, R.S. and Jones, P.D. (eds.) Climate Since A.D. 1500, pp 118-142, Routledge, London, New York. Plaut, G., Ghil, M. and Vautard, R. (1995): Interannual and Interdecadal Variability in 335 Years of Central England Temperatures. Science 268, 710-713. PTT, G. (1960): Malojastrasse Oberengadin-Bergell. Generaldirektion der Post-, Telegraphen- und Telephonverwaltung, 62p, Bern. Raymo, M.E., Ruddiman, W.F., Shackleton, N. and Oppo, D.W. (1990): Evolution of Atlantic-Pacific ∂13C gradients over the last 2.5 m.y. Earth and Planetary Science Letters 97, 353-368. Reid, G.C. (1997): Solar Forcing of Global Climate Change Since The Mid-17th Century. Climatic Change 37(2), 391-405. Renberg, I., Segerström, U. and Wallin, J.-E. (1984): Climatic Reflection in Varved Lake Sediments. In: Mörner, N.-A. and Karlén, W. (eds.) Climatic Changes on a yearly to Millennial Basis, pp 249-256, D. Reidel Publishing Company, Dordrecht, Holland. Robbi, J. (1913): Die Heilquellen von St. Moritz bis 1815. - historische Skizze von Dr. Jules Robbi. Selbstverlag, Chur. Robinson, A.B. and Robinson, Z.W. (1997): Science Has Spoken: Global Warming Is a Myth. The Wall Street Journal Europe Nov. 5-6, 10. Röthlisberger, G. (1991): Chronik der Unwetterschäden in der Schweiz. Eidgenössische Forschungsanstalt für Wald, Schnee und Landschaft, 122p, Birmensdorf, Schweiz. Rupke, N.A. (1978): Deep Clastic Seas. In: Reading, H.G. (eds.) Sedimentary Environments and Facies, pp 372-415, Blackwell Scientific Publications, Oxford, London, etc.. SANW. (1986-1994): Die Gletscher der Schweizer Alpen. Jahrbuch der Gletscherkommission der schweizerischen Akademie der Naturwissenschaften 99-112, Versuchsanstalt für Wasserbau, Hydrologie und Glaziologie, Zürich. Sarnthein, M., Thiede, J., Pflaumann, U., Erlenkeuser, H., Fütterer, D., Koopmann, B., Lange, H. and Seibold, E. (1982): Atmospheric and oceanic circulation patterns off northwest Africa during the past 25 million years. In: von Rad, U., Sarnthein, M. and Seibold, E. (eds.) Geology of the Northwest African Continental Margin, pp 584-604, Springer-Verlag, New York. References 185 References

Schaaf, M. and Thurow, J. (1994): A fast and easy method to derive highest-resolution time-series datasets from drillcores and roch samples. Sedimentary Geology 94, 1-10. Schaaf, M. and Thurow, J. (1995): Late Pleistocene-Holocene climatic cycles recorded in Santa Barbara Basin sediments: Interpretation of Color density logs from site 893. In: Kennett, J.P., Baldauf, J.G. and Lyle, M. (eds.) Proceedings of the Ocean Drilling Program, Scientific Results (Part 2), Vol. 146, pp 31-44, Ocean Drilling Program, Texas A&M University. Schaller, T.L. (1996): Redox-Sensitive Metals in Recent Lake Sediments Proxy-Indicators of Deep- Water Oxygen and Climate Conditions. Ph.D. thesis No. 11901, ETH Zürich. Schlüchter, C. (1988): The Deglaciation of the Swiss-Alps: A Paleoclimatic Event with Chronological Problems. Bulletin de l’Association française pour l’étude du Quaternaire 2/3, 141-145. Schmassmann, W. (1920): Die Bodenfauna hochalpiner Seen. Archiv für Hydrobiologie Supplement III(1), 1-106. Schmid, E. (1978): Vegetationskarte der Schweiz (1:500000). In: Imhof, E. (ed.) Atlas der Schweiz, Eidgenössische Landestopographie, Wabern-Bern. Schmid, S.M. and Froitzheim, N. (1993): Oblique slip and block rotation along the Engadine line. Eclogae Geologicae Helvetiae 86(2), 569-593. Schneider, S.H. (1990): Wissenschaftliche Klimamodelle - Ein Beitrag zur Diskussion über globale Erwärmung. In: Leggett, J. (eds.) Global Warming, pp 74-105, R. Piper GmbH & Co.KG, München. Schoeneich, P. (1997): Les stades tardiglaciares des Préalpes vaudoises et leur corrélation avec les modèle des Alpes orientales. Jahresversammlung Schweizerische Geomorphologische Gesellschaft, 1. Schönwiese, C.-D. (1994): Klimatologie. Ulmer Verlag, 436p, Stuttgart. Schüepp, M. (1967): Klimatologie der Schweiz, Lufttemperatur. Beiheft zu den Annalen der Scheizerischen Meteorologischen Zentralanstalt 3/4, 63-106. Schweingruber, F.H. (1987b): Flächenhafte dendroklimatische Temperaturrekonstruktionen für Europa. Naturwissenschaften 74, 205-212. Schweingruber, F.H. (1989): Tree Rings. Kluwer Academic Publishers, 276p, Dordrecht Boston, London. Schwoerbel, J. (1980): Einführung in die Limnologie. Gustav Fischer Verlag, 196p, Stuttgart, New York. Seger, F. (1971): Geologische Untersuchungen in der Fexer Schuppenzone zwischen Sils-Maria und Fex-Curtins. Diploma-Thesis, 128p, ETH Zürich.

Sharp, M., Tranter, M., Brown, G.H. and Skidmore, M. (1995): Rates of chemical denudation and CO2 drawdown in a glacier-covered alpine catchment. Geology 23(1), 61-64. Siegenthaler, C. and Sturm, M. (1991): Die Häufigkeit von Ablagerungen extremer Reuss-Hochwasser. Die Sedimentationsgeschichte im Urnersee seit dem Mittelalter. Mitteilungen des Bundesamtes für Wasserwirtschaft 4, 127-139. Sils, J. (1993/94): Jahresbericht Kur- und Verkehrsverein Sils, Engadin. Silvaplana, J. (1993/94): Jahresbericht Kur- und Verkehrsverein Silvaplana-Surlej-Champfèr, Engadin. SMA. (1864-1994): Annalen der Schweizerischen Meteorologischen Anstalt, Zürich. Small, R.J. (1987a): Englacial and Supraglacial Sediment: Transport and Deposition. In: Gurnell, A.M. and Clark, M.J. (eds.) Glacio-fluvial sediment transfer. An Alpine Perspective, pp 111-145, Wiley & Sons, Chichester, New York, Brisbane. Small, R.J. (1987b): Moraine sediment budgets. In: Gurnell, A.M. and Clark, M.J. (eds.) Glacio-fluvial sediment transfer. An Alpine Perspective, pp 165-197, Wiley & Sons, Chichester, New York, Brisbane. Smith, D.G. (1994): Cyclicity or chaos? Orbital forcing versus non-linear dynamics. Spec. Publs. Int. Ass. Sediment. 19, 531-544. References 186 References

Snowball, I. (1993): Mineral magnetic properties of Holocene lake sediments and soils from the Kårsa valley, Lappland, Sweden, and their relevance to palaeoenvironmental reconstruction. Terra Nova 5, 258-270. Souchez, R.A. and Lorrain, R.D. (1987): The Subglacial Sediment System. In: Gurnell, A.M. and Clark, M.J. (eds.) Glacio-fluvial sediment transfer. An Alpine Perspective, pp 147-164, Wiley & Sons, Chichester, New York, Brisbane. Spillmann, P. (1993): Die Geologie des penninisch-ostalpinen Grenzbereichs im südlichen Berninagebirge. PhD Thesis No. 10175, ETH Zürich. Spillmann, P. and Büchi, H.J. (1993): The Pre-Alpine Basement of the Lower Austro-Alpine Nappes in the Bernina Massif (Grisons, Switzerland; Valtellina, Italy). In: Raumer, J.F. and Neubauer, F. (eds.) Pre-Mesozoic Geology in the Alps, pp 457-467, Springer-Verlag, Berlin Heidelberg. St. Arnaud, R.J. and Sudom, M.D. (1981): Mineral Distribution and Weathering in Chernozemic and Luvisolic Soild from Central Saskatchewan. Canadian Journal of Soil Science 61, 79-89. Stallard, R.F. and Edmond, J.M. (1983): Geochemistry of the Amazon 2. The influence of geology and weathering environment on the dissolved load. Journal of geophysical research 88(C14), 9671-9688. Staub, J. (1864): Die Pfahlbauten in den Schweizer Seen. Verlag des Verfassers, Fluntern bei Zürich. Staub, R. (1938): Zur Frage einer Schlussvereisung im Berninagebiet zwischen Bergell, Oberengadin und Puschlav. Eclogae Geologicae Helvetiae 31(1), 125-136. Staub, R. (1946): Geologische Karte der Bernina-Gruppe und Umgebung, Geologische Spezialkarte der Schweiz, Blatt 118 (1:50000). Staub, R. (1952): Der Paß von Maloja. Seine Geschichte und Gestaltung. Separatabdruck aus dem Jahresbericht der Naturforschenden Gesellschaft Graubündens 83, 84. Stihler, S.D., Stone, D.B. and Beget, J.E. (1992): “Varve” counting vs. tephrochronology and 137Cs and 210Pb dating: A comparative test at Skilak Lake, Alaska. Geol 20, 1019-1022.

14 Stuiver, M. and Braziunas, T.F. (1993): Sun, ocean, climate and atmospheric CO2: an evaluation of causal and spectral relationships. The Holocene 3(4), 289-305. Stuiver, M. and Reimer, P.J. (1993): Extended 14C database and revised CALIB radiocarbon calibration program. Radiocarbon 35, 215-230. Stumm, W. and Morgan, J.J. (1981): Aquatic Chemistry. John Wiley & Sons, 780p, NewYork, Chichester, etc. Sturm, M. (1979): Origin and composition of clastic varves. In: Schlüchter, C. (eds.) Moraines and Varves, pp 281-285, A.A. Balkema, Rotterdam. Sturm, M. and Matter, A. (1978): Turbidites and varves in Lake Brienz (Switzerland): deposition of clastic detritus by density currents. Special Publications international Association Sediment. 2, 147- 168. Suter, J. (1981): Gletschergeschicte des Oberengadins: Untersuchung von Gletscherschwankungen in der Err-Julier- Gruppe. Ph.D. Thesis Univ. Zürich, Physische Geographie 2. Suter, J. (1991): Die Tiefbohrung Sils im Oberengadin im Lichte der Glazialmorphologie. In: Elsasser, H. and Boesch, M. (eds.) Beiträge zur Geographie Graubündens, pp 46-49, Fotorotar AG, Zurich. Suter, J., Gamper-Schollenberger, B., Zoller, H., Heitz-Weniger, A. and Punchakunnel, P. (1982): Gletscher, Vegetations und Klimageschichte im Oberengadin. In: Maisch, M. and Suter, J. (eds.) Physische Geographie. Exkursionsfuhrer: Teil A. Ostschweiz, Vol. 6, pp 14-30, Zurich. Sweeney, R.E. and Kaplan, I.R. (1973): Pyrite Framboid Formation: Laboratory Synthesis and Marine Sediments. Econ. Geology 68, 618-634. Teranes, J.L. and McKenzie, J.A. (1995): Evidence for rapid climate changes during the 20th century from high-resolution oxygen isotope stratigraphy in chemically varved lacustrine sediments. Terra nova 7(Abstract supplement No. 1), 218. References 187 References

Thomas, R.L. (1988): Lake sediments as indicators of changes in land erosion rates. In: Lerman, A. and Meybeck, M. (eds.) Physical and Chemical Weathering in Geochemical Cycles, pp 143-164, Kluver Academic Publishers, . Thompson, R., Battarbee, R.W., O’Sullivan, P.E. and Oldfield, F. (1975): Magnetic susceptibility of lake sediments. Limnology and Oceanography 20(5), 687-698. Tiedemann, R., Sarnthein, M. and Shackleton, N.J. (1994): Astronomic timescale for the Pliocene Atlantic ∂18O and dust flux records of Ocean Drilling Program site 659. Paleoceanography 9(4), 619-638. Trümpy, R. (1977): The Engadine Line: a sinsistral wrench fault in the Central Alps. Mem. Geol. Soc. China 2, 1-17. Trümpy, R. (1980): Geology of Switzerland, a guide-book. Wepf & Co. Publishers, 334p, Basel, New York. Turekian, K.K. and Wedepohl, K.H. (1961): Distribution of the Elements in Some Major Units of the Earth’s Crust. Geological Society of America Bulletin 72, 175-192. Udden, J.A. (1914): Mechanical Composition of Clastic Sediments. Bull. Geol. Soc. Am. 25, 655-744. Urfer, C., Gensler, G., Ambrosetti, F. and Zenone, E. (1979): Klimatologie der Schweiz. Regionale Klimabeschreibungen. 2. Teil Alpennordhang, Graubünden und Alpensüdseite. Beiheft zu den Annalen der Scheizerischen Meteorologischen Anstalt 19(2), 247-437. Uttinger, H. (1966): Klimatologie der Schweiz, Niederschlag. Beiheft zu den Annalen der Scheizerischen Meteorologischen Zentralanstalt 4, 125-170. Vali, H., Förster, O., Amarantidis, G. and Petersen, N. (1987): Magnetotactic bacteria and their magnetofossils in sediments. Earth and Planetary Science Letters 86, 389-400. van Husen, D. (1997): LGM and Late-Glacial Fluctuations in the eastern Alps. Quaternary International 38/39, 109-118. Velbel, M.A. (1993): Temperature dependance of silicate weathering in nature: How strong a negative

feedback on long term accumulation of atmospheric CO2 and global greenhouse warming ? Geology 21, 1059-1062. Venetz, I. (1833): Mémoire sur les variations de la température dans les Alps de la Suisse. Denkschrift der schweizerischen Gesellschaft für die gesamten Naturwissenschaften 1(2). Vogel, H.-H. (1995): Geomorphologische Kartierung sowie Gletschergeschchtliche und Geochronologische Untersuchungen im Val Fedoz. Diploma thesis, 95p, Geographical Institute University of Zürich. von Blanckenburg, F. (1992): Combined high-precision chronometry and geochemical tracing using accessory minerals; applied to the Central-Alpine Bergell intrusion. Chemical Geology 100, 19-40. Walker, M.J.C. (1995): Climatic Changes in Europe during the Last Glacial/Interglacial Transition. Quaternary International 28, 63-76. Walling, D.E. (1987): Rainfall, Runoff and Erosion of the Land: A Global View. In: Gregory, K.J. (eds.) Energetics of Physical Environment, pp 89-117, John Wiley & Sons Ltd.. Walling, D.E., Peart, M.R., Oldfield, F. and Thompson, R. (1979): Suspended sediment sources identified by magnetic maesurements. Nature 281, 110-113. Walter, H. (1990): Vegetation und Klimazonen. Grundriss der globalen Oekologie. Verlag Eugen Ulmer, 382p, Stuttgart. Wanner, W.J. (1971): Zur Geologie der Fexer Schuppenzone zwischen Fex-Curtins und Piz Tremoggia. Diploma-Thesis, 149p, ETH Zürich. Wasserwirtschaft, B.f. (1991): Ursachenanalyse der Hochwasser 1987; Schlussbericht. EDMZ, 47p, Bern. Wedepohl, K.H. (1978): Handbook of Geochemistry. Springer-Verlag, 30 B-Op, Berlin, Heidelberg, New York. References 188 References

Weissert, H.J. and Bernoulli, D. (1985): A transform margin in the Mesozoic Tethys: evidence from the Swiss Alps. Geologische Rundschau 74(3), 665-679. Welch, P.D. (1967): The Use of Fast Fourier Transform for the Estimation of Power Spectra: A Method Based on Time Averaging Over Short, Modified Periodograms. IEEE Transactions on Audio and Electroacoustics AU-15(2), 70-73. Wendler, G. and Weller, G. (1974): A Heat-Balance study on McCall Glacier, Brooks Range, Alaska: A contribution to the International Hydrological Decade. Journal of Glaciology 13(67), 13-26. Wentworth, C.K. (1922): A Scale of Grade and Class Terms for Clastic Sediments. J. Geol. 30/5, 377-392. Wessels, M. (1995): Bodensee-Sedimente als Abbild von Umweltänderungen im Spät- und Postglazial. Göttinger Arbeiten zur Geologie und Paläontologie 66, 105p. Wohlfarth, B., Gaillard, M.-J., Haeberli, W. and Kelts, K. (1994): Environment and Climate in Southwestern Switzerland during the last Termination, 10-15 ka BP. Quaternary Science Reviews 13, 361-394. Worm, H.-U., Nowaczyk, N.R. and Scheffler, M. (1995): Application of the HFD: Magnetic Field Scanning of Varved Lake Sediments. Spot Beam, information of the Forschungsgesellschaft für Informationstechnik mbH 9(special Issue/HFD), 1-3. Ziegler, J.M. (1876): Ueber das Verhältniss der Topographie zur Geologie. Text zur topographischen Karte vom Engadin und Bernina. Ein Beitrag zur Geschichte der Erdkruste, Verlag von J. Wurster & Comp., Zürich. Zoller, H. (1977): Alter und Ausmass postglazialer Klimaschwankungen in den Schweizer Alpen. In: Frenzel, B. (eds.) Dendrochronologie und Postglaziale Klimaschwankungen in Europa, pp 271-281, Franz Steiner Verlag GmbH, Wiesbaden. Zoller, H., Erny-Rodmann, C. and Punchakunnel, P. (1996): The history of vegetation and land use in the Lower Engadine (Switzerland). Pollen record of the last 13000 years. Nationalpark-Forschung in der Schweiz 86, 65p. Zuan, H. (1984): Sils i. E. einst und jetzt Gemeindeportrait von Hnas Zuan. Buchdruckerei St. Moritz AG, St Moritz. Züllig, H. (1982): Die Entwicklung von St. Moritz zum Kurort im Spiegel der Sedimente des St. Moritzersees. Wasser, Energie, Luft 74(7/8), 177-183. Zumbühl, H. and Messerli, P. (1980): Gletscherschwankungen und Temperaturverlauf. Beispiel einer Korrelationsanalyse von indirekten und direkten Klimazeugen am Beispiel der Grindelwaldgletscher und er 210-jährigen Baseler Temperaturreihe. In: Oeschger, H., Messerli, B. and Svilar, M. (eds.) Das Klima. Analysen und Modelle Geschichte und Zukunft, pp 161-174, Springer- Verlag, Berlin, etc... Appendices

Appendices 191 Appendices

Appendix 5.1: Minerals detected by XRD in the surface sediments of Lej da Segl (SS) Lej da Silvaplauna (SP) and Lej da Champfèr (CH). Peak intensities for characteristic peaks are given. Sample Sed.depth Chlorite Mica Talc Serp. Amph. Quartz Quartz Kfsp. Plag Dol. Calcite Cpx No. Core No. [cm] [7.00Å] [10.00Å] [9.33Å] [7.30Å] [8.45Å] [4.24Å] [3.34Å] [3.24Å] [3.19Å] [2.89Å] [3.03Å] [2.99Å] Ofl01 SS 91/1 0-1 399 377 18 66 86 362 31 154 49 Stilp. [42] Ofl02 SS 91/2 0-1 400 412 120 156 1035 89 291 42 70 Stilp. [36] Ofl03 SS 91/3 0-1 212 265 119 173 642 38 235 53 53 Stilp. [24] Ofl04 SS 91/4-1 0-1 430 318 102 121 771 142 307 36 57 Stilp. [28] Ofl05 SS 91/5 0-1 319 286 81 75 438 36 111 31 34 Stilp. [22] Ofl06 SS 92/2-1 0-1 272 181 22 54 115 636 121 479 38 57 Ofl07 SS 92/3-1 0-1 379 227 16 78 182 768 782 411 55 30 30 Ofl08 SS 92/5-1 0-1 415 448 90 155 726 49 245 31 63 Ofl09 SS 92/6-1 0-1 238 282 112 147 791 54 260 37 69 Ofl10 SS 93/1 0-1 278 199 16 49 108 473 35 142 35 23 36 Ofl11 SS 93/2 0-1 211 194 26 45 82 393 103 31 61 Ofl12 SS 93/3 0-1 213 218 45 56 324 181 48 Ofl13 SS 93/4 0-1 383 396 83 250 1363 41 608 53 87 Ofl14 SS 91/1 4.0-5.0 134 125 74 131 769 47 190 45 84 Ofl15 SS 91/2 7.0-8.0 385 493 77 310 1248 68 375 62 82 Ofl16 SS 91/2 12.5-13.5 247 274 157 172 789 76 257 59 70 Ofl17 SS 91/3 6.0-7.0 199 237 97 230 922 66 201 55 84 Ofl18 SS 91/4-1 15.0-16.0 490 276 166 118 622 52 245 50 80 Ofl19 SS 91/4-2 30.0-31.0 1149 988 96 172 945 58 288 55 101 Ofl20 SS 92/6-1 16.0-17.0 202 330 123 412 1419 62 346 97 86 Ofl21 SS 92/6-1 17.0-18.0 324 434 161 199 1089 86 283 48 58 Ofl22 SS 91/5 10.0-11.0 418 253 35 144 131 521 31 160 47 79 Ofl23 SS 92/2-1 7.0-8.0 276 157 75 84 324 39 144 28 45 Ofl24 SS 92/2-1 11.0-12.0 373 264 38 153 105 457 40 133 60 73 Stilp. [23] Ofl25 SS 92/3-1 16.0-17.0 498 310 126 134 522 80 39 54 Ofl26 SS 92/3-1 34.0-34.5 170 298 89 188 862 57 207 28 68 Ofl27 SS 92/5-1 7.5-8.5 449 415 122 136 766 49 176 38 78 Stilp. [21] Ofl28 SS 93/1 4.5-5.5 209 281 32 242 1172 35 211 321 67 Ofl29 SS 93/1 10.0-11.0 542 293 29 49 145 115 517 39 125 42 58 Ofl30 SS 93/2 2.5-3.0 858 555 65 39 69 147 642 87 Ofl31 SS 93/2 11.0-12.0 387 299 80 87 390 37 33 60 Ofl32 SS 93/3 8.0-9.0 412 299 76 88 468 31 138 47 Stilp. [24] Ofl33 SS 93/4 10.0-11.0 350 388 26 135 139 837 93 245 43 77 Ofl34 SSP 93/1 0-1 712 604 32 47 208 1141 37 321 84 Ofl35 SSP 93/2 0-1 698 1091 75 341 1509 80 1286 103 72 Ofl36 SSP 93/4 0-1 584 613 45 125 686 29 191 78 Ofl37 SSP 93/5 0-1 749 664 56 198 787 560 81 84 Ofl38 PSP 90/2-0 0-1 180 218 41 184 999 46 260 109 61 Ofl39 PSP 90/3-1 0-1 262 307 39 136 771 44 261 50 64 Stilp. [22] Ofl40 SSP 92/1 0-1 302 293 55 278 1144 45 254 100 68 Ofl41 SSP 92/4-1 0-1 432 372 54 182 890 54 292 110 61 Ofl42 SSP 92/6-1 0-1 219 248 37 232 1021 46 257 109 65 Ofl43 SSP 92/7-1 0-1 346 298 30 51 169 989 67 228 81 60 Ofl44 SSP 93/1 38.0-39.0 794 978 66 247 1350 36 465 84 124 Ofl45 SSP 93/2 10.0-11.0 528 610 64 290 1337 38 391 91 77 Ofl46 SSP 93/4 10.0-11.0 912 901 71 188 1109 41 593 31 109 Ofl47 SSP 93/5 26.0-27.0 1192 1258 34 79 296 1298 75 347 322 91 Ofl48 SSP 93/5 53.0-53.5 431 647 69 306 1512 68 372 62 83 Ofl49 PSP 90/2-0 17.0-18.0 619 590 30 56 191 886 41 1070 99 83 Ofl50 PSP 90/3-1 16.0-17.0 607 500 31 42 172 1076 45 380 64 81 Ofl51 PSP 90/3-1 80.5-81.5 874 1263 67 339 1554 453 469 63 113 Ofl52 SSP 92/1 16.0-17.0 905 760 29 86 180 963 51 394 51 74 Ofl53 SSP 92/1 30.0-31.0 846 1081 129 68 227 1151 40 545 140 87 Ofl54 SSP 92/4-1 20.0-21.0 567 500 70 164 872 33 346 77 Stilp. [29] Ofl55 SSP 92/7-1 23.0-24.0 537 485 49 239 1165 73 427 34 93 Ofl56 SSP 92/6-1 17.0-18.0 582 576 21 58 237 1185 81 286 247 86 Ofl57 CF 92/2-2-1 0-1 264 271 115 398 102 56 52 Ofl58 CF 93/2 0-1 113 120 54 162 239 27 Ofl59 CF 93/1 0-1 195 189 26 90 359 117 38 43 Ofl60 CF 92/2-2-1 21.0-22.0 216 197 29 87 436 37 152 34 58 Ofl61 CF 92/2-2-1 60.5-61.5 163 189 179 915 68 254 38 44 Ofl62 CF 93/2 26.0-27.0 197 151 28 123 687 63 235 69 Ofl63 CF 93/2 39.0-40.0 544 450 33 90 501 205 88 Ofl64 CF 93/1 20.0-21.0 443 357 98 612 30 233 81

Appendix 5.2: Average element accumulation rates determined with sedimentation rates given in Leemann (1993) for mix and spot samples from core PSP 90/2. MAR Na Ca Mg Al Fe Name Core No. Age (yBP) [g/cm2*yr] [mg/cm2*yr] K [mg/cm2*yr] [mg/cm2*yr] [mg/cm2*yr] [mg/cm2*yr] [mg/cm2*yr] Fe/Mn 1mix SSP 92/6-1 0-500 0,25 3,15 7,32 4,79 6,52 17,93 12,75 4871,03 277,78 239,83 41,17 26,17 2amix PSP 90/1-1,2 500-1800 0,22 2,95 8,03 3,46 6,60 19,83 12,89 3493,24 283,49 180,08 28,24 36,90 2bmix PSP 90/1-3,4,5,6 1500-3800 0,17 2,28 6,28 2,59 5,11 15,41 10,16 2947,91 213,69 276,84 24,89 34,46 3mix PSP 90/2-10,11,12 3800-9415 0,025 0,34 0,76 0,57 0,82 2,02 1,28 447,19 33,24 22,26 3,03 28,71 4mix PSP 90/2-12 >9415 BP 0,12 1,82 4,34 2,03 4,39 10,16 6,07 1683,24 137,69 97,92 17,69 36,06 PSP 90/3 recent basin 0,28 4,16 10,94 5,35 8,98 31,00 17,82 9161,89 433,48 297,33 12,38 19,45 1spot PSP 90/2 recent slope 0,25 3,94 9,52 6,21 8,39 26,36 14,50 3621,18 364,03 227,58 19,08 40,05 2spot PSP 90/1-4 ca. 1800 BP 0,17 2,13 6,59 2,17 5,21 16,32 10,93 2371,67 213,50 314,00 18,67 46,10 3spot PSP 90/1-8 ca. 6000 BP 0,025 0,29 0,75 0,56 0,71 1,92 1,24 756,10 39,26 23,75 2,17 16,40 Appendices 192 Appendices

Appendix 6.1: MAR and contents of water, BiSi, C_org, C_carb, total sulfur, total nitrogen and total hydrogen measured in samples taken from core PSP 94/3-0. For calculation of accumulation rates gaps, where no measurements were conducted, were filled by linear interpolation between the two nearest datapoints. MAR were calculated from the varve thickness measuremsnts.

Total Total Water Biogenic organic inorganic Total Total Total Sample MAR content silica carbon carbon sulfur nitrogen hydrogen Year No. [mg/cm2*yr] [%] [%] [%] [%] [%] [%] [%] 1994 1 108 72,6 11,52 2,84 0,14 0,00 0,34 0,98 1993 2 232 59,0 3,45 1,11 0,14 0,19 0,11 0,68 1992 3 147 70,1 6,39 1,87 0,37 0,00 0,20 0,76 1991 4 75 74,7 9,79 2,25 0,14 0,51 0,23 0,90 1990 85 1989 5 67 75,8 10,64 2,43 0,09 0,59 0,27 0,97 1988 6 90 70,3 7,83 1,89 0,10 0,00 0,21 0,88 1987 59 1986 10 100 76,9 14,67 2,50 0,29 0,75 0,29 0,91 1985 11 104 69,6 7,59 1,96 0,35 0,39 0,20 0,80 1984 12 39 76,8 10,77 2,73 0,31 1,22 0,30 0,90 1983 13 119 70,4 11,27 2,12 0,20 0,60 0,22 0,80 1982 14 155 73,1 10,73 1,87 0,16 0,49 0,21 0,80 1981 15 126 71,5 9,19 1,88 0,24 0,54 0,22 0,82 1980 16 88 73,2 10,04 1,91 0,20 0,37 0,21 0,83 1979 17 77 69,6 7,10 2,35 0,26 0,34 0,26 0,87 1978 18 137 63,0 3,96 1,54 0,28 0,36 0,15 0,69 1977 19 85 76,3 16,66 2,05 0,29 0,99 0,21 0,77 1976 69 1975 20 81 81,9 17,72 2,78 0,16 1,25 0,31 0,93 1974 21 112 77,1 11,77 2,14 0,19 0,82 0,30 0,92 1973 23 56 78,2 12,82 2,39 0,29 0,98 0,32 0,98 1972 24 42 81,0 15,15 2,63 0,29 1,11 0,37 1,07 1971 25 48 80,9 13,36 2,22 0,32 1,02 0,32 0,95 1970 35 1969 26 65 77,0 11,57 2,13 0,32 0,78 0,29 0,91 1968 37 1967 27 61 76,9 13,50 2,40 0,42 0,95 0,32 0,92 1966 66 1965 28 45 73,3 12,78 2,28 0,22 1,00 0,30 1,08 1964 29 70 70,3 9,30 1,83 0,32 0,76 0,23 0,94 1963 81 1962 30 47 70,3 9,21 1,55 0,51 0,83 0,21 0,80 1961 31 59 69,1 6,92 1,65 0,26 0,65 0,22 0,81 1960 90 1959 33 101 61,3 3,62 1,13 0,38 0,32 0,15 0,66 1958 102 1957 59 1956 38 103 47,8 1,26 0,56 0,41 0,00 0,06 0,54 1955 158 1954 39 70 50,2 1,91 0,74 0,65 0,00 0,09 0,61 1953 170 1952 40 96 42,8 1,19 0,48 0,39 0,00 0,05 0,52 1951 197 1950 41 323 44,8 1,36 0,58 0,24 0,00 0,06 0,56 1949 42 264 45,1 1,21 0,42 0,32 0,00 0,05 0,60 1948 43 172 47,9 1,41 0,58 0,64 0,00 0,06 0,57 1947 44 342 44,0 1,10 0,43 0,93 0,00 0,05 0,54 1946 228 1945 46 122 46,2 1,20 0,61 1,09 0,00 0,06 0,52 1944 218 1943 232 1942 47 165 52,9 2,03 0,90 0,07 0,00 0,11 0,67 1941 48 96 48,0 1,41 0,65 0,11 0,00 0,07 0,59 1940 114 1939 85 1938 163 1937 49 151 45,5 1,30 0,50 0,07 0,00 0,07 0,54 1936 167 1935 50 99 47,1 1,79 0,50 0,10 0,00 0,07 0,55 1934 199 1933 103 1932 108 Appendices 193 Appendices

Appendix 6.1: continuation

Total Total Water Biogenic organic inorganic Total Total Total Sample MAR content silica carbon carbon sulfur nitrogen hydrogen Year No. [mg/cm2*yr] [%] [%] [%] [%] [%] [%] [%] 1931 51 120 44,8 1,31 0,42 0,12 0,00 0,06 0,54 1930 168 1929 52 128 40,6 0,86 0,41 0,23 0,00 0,04 0,51 1928 53 129 43,6 0,92 0,52 0,21 0,00 0,06 0,57 1927 117 1926 54 116 49,2 1,51 0,80 0,16 0,00 0,10 0,61 1925 81 1924 55 123 48,8 1,79 0,66 0,03 0,00 0,08 0,63 1923 127 1922 144 1921 56 246 45,6 1,43 0,64 0,08 0,00 0,07 0,59 1920 93 1919 127 1918 57 108 49,1 1,83 0,78 0,15 0,00 0,10 0,65 1917 140 1916 99 1915 58 101 50,0 1,86 0,93 0,09 0,00 0,12 0,66 1914 36 1913 76 1912 68 1911 83 1910 76 1909 77 1908 61 113 46,2 1,48 0,67 0,06 0,00 0,09 0,62 1907 139 1906 62 154 45,4 1,33 0,43 0,05 0,00 0,06 0,56 1905 171 1904 63 196 43,9 1,32 0,41 0,06 0,00 0,06 0,57 1903 120 1902 97 1901 64 189 40,6 1,05 0,36 0,06 0,00 0,06 0,56 1900 202 1899 172 1898 65 149 35,9 1,14 0,42 0,06 0,00 0,07 0,59 1897 151 1896 216 1895 213 1894 66 268 37,6 0,87 0,33 0,11 0,00 0,05 0,54 1893 185 1892 67 232 32,9 1,08 0,38 0,04 0,00 0,06 0,60 1891 228 1890 200 1889 68 386 35,7 2,17 0,48 0,10 0,00 0,08 0,64 1888 330 1887 72 149 42,6 1,10 0,75 0,11 0,00 0,09 0,65 1886 105 1885 93 1884 73 273 37,4 1,06 0,42 0,09 0,00 0,06 0,58 1883 110 1882 74 103 40,9 1,13 0,51 0,08 0,00 0,07 0,63 1881 157 1880 73 1879 76 104 34,7 0,80 0,37 0,11 0,00 0,06 0,56 1878 126 1877 115 1876 223 1875 77 91 39,5 0,89 0,37 0,07 0,00 0,06 0,59 1874 317 1873 120 1872 78 251 39,0 0,97 0,39 0,05 0,00 0,06 0,61 1871 279 1870 79 210 31,4 0,89 0,39 0,11 0,00 0,06 0,56 1869 181 Appendices 194 Appendices

Appendix 6.1: continuation

Total Total Water Biogenic organic inorganic Total Total Total Sample MAR content silica carbon carbon sulfur nitrogen hydrogen Year No. [mg/cm2*yr] [%] [%] [%] [%] [%] [%] [%] 1868 179 1867 122 1866 80 260 40,2 0,83 0,37 0,09 0,00 0,06 0,60 1865 224 1864 82 256 33,6 0,90 0,35 0,14 0,00 0,06 0,58 1863 201 1862 118 1861 172 1860 85 187 37,9 0,66 0,33 0,11 0,00 0,05 0,58 1859 247 1858 132 1857 232 1856 86 138 39,0 0,77 0,34 0,05 0,00 0,06 0,59 1855 87 230 31,1 0,75 0,32 0,11 0,00 0,06 0,58 1854 244 1853 207 1852 88 150 35,4 0,96 0,35 0,09 0,00 0,07 0,60 1851 255 1850 303 1849 274 1848 113 1847 127 1846 90 334 39,9 0,72 0,37 0,12 0,00 0,09 0,58 1845 91 269 40,4 0,74 0,34 0,07 0,00 0,01 0,55 1844 92 277 41,4 0,96 0,35 0,05 0,00 0,02 0,57 1843 93 296 40,4 0,66 0,34 0,10 0,00 0,01 0,54 1842 94 329 40,8 0,80 0,40 0,12 0,00 0,01 0,56 1841 95 288 40,8 0,87 0,33 0,09 0,00 0,02 0,56 1840 96 299 40,1 0,90 0,36 0,10 0,00 0,01 0,56 1839 97 296 40,9 0,70 0,39 0,08 0,00 0,02 0,56 1838 92 1837 213 1836 99 193 39,9 0,68 0,42 0,19 0,00 0,02 0,56 1835 100 324 38,6 0,71 0,34 0,23 0,00 0,02 0,56 1834 94 1833 101 368 37,8 0,82 0,34 0,17 0,00 0,02 0,56 1832 70 1831 102 365 36,5 0,77 0,33 0,20 0,00 0,01 0,53 1830 103 81 40,6 0,81 0,41 0,17 0,00 0,02 0,59 1829 150 1828 179 1827 104 126 40,2 0,95 0,40 0,10 0,00 0,02 0,57 1826 196 1825 209 1824 105 146 41,5 0,97 0,45 0,13 0,00 0,03 0,60 1823 175 1822 173 1821 107 236 38,6 0,65 0,39 0,29 0,00 0,02 0,57 1820 237 1819 108 143 40,8 0,94 0,43 0,31 0,00 0,03 0,61 1818 162 1817 243 1816 109 174 42,8 0,92 0,37 0,06 0,00 0,02 0,58 1815 166 1814 130 1813 110 156 38,8 1,00 0,33 0,08 0,00 0,02 0,57 1812 202 1811 354 1810 238 1809 145 1808 143 1807 107 1806 111 155 41,4 0,75 0,35 0,07 0,00 0,01 0,56 Appendices 195 Appendices

Appendix 6.1: continuation

Total Total Water Biogenic organic inorganic Total Total Total Sample MAR content silica carbon carbon sulfur nitrogen hydrogen Year No. [mg/cm2*yr] [%] [%] [%] [%] [%] [%] [%] 1805 245 1804 189 1803 217 1802 112 178 41,4 0,95 0,38 0,09 0,00 0,02 0,56 1801 237 1800 107 1799 304 1798 222 1797 340 1796 113 352 42,6 0,87 0,36 0,05 0,00 0,02 0,61 1795 252 1794 226 1793 114 88 47,0 1,37 0,52 0,02 0,00 0,04 0,62 1792 121 1791 118 1790 65 1789 115 145 43,6 1,12 0,44 0,01 0,00 0,03 0,58 1788 206 1787 114 1786 131 1785 136 1784 117 1783 262 1782 207 1781 116 102 43,2 1,12 0,42 0,02 0,00 0,03 0,59 1780 202 1779 130 1778 232 1777 157 1776 116a 126 0,02 1775 110 1774 58 1773 191 1772 116b 115 0,03 1771 163 1770 171 1769 116c 105 0,14 1768 219 1767 116d 91 0,24 1766 87 1765 137 1764 142 1763 132 1762 55 1761 46 1760 50 1759 116e 59 0,30 1758 46 1757 50 1756 46 1755 100 1754 119 1753 96 1752 146 1751 69 1750 116f 64 0,24 0,22 0,01 1749 59 1748 55 1747 41 1746 114 1745 91 1744 87 1743 78 Appendices 196 Appendices

Appendix 6.1: continuation

Total Total Water Biogenic organic inorganic Total Total Total Sample MAR content silica carbon carbon sulfur nitrogen hydrogen Year No. [mg/cm2*yr] [%] [%] [%] [%] [%] [%] [%] 1742 69 1741 50 1740 69 1739 64 1738 142 1737 116g 114 0,27 0,17 0,01 1736 187 1735 265 1734 82 1733 119 1732 116h 155 0,75 0,12 0,05 1731 174 1730 116i 178 0,52 0,14 0,02 1729 91 1728 146 1727 192 1726 215 1725 169 1724 128 1723 237 1722 219 1721 123 1720 160 1719 183 1718 146 1717 132 1716 131 1715 117 166 41,8 0,95 0,44 0,06 0,00 0,03 0,58 1714 126 1713 135 1712 131 1711 198 1710 118 147 40,5 1,05 0,49 0,19 0,00 0,04 0,57 1709 115 1708 118 1707 192 1706 99 1705 123 1704 135 1703 165 1702 201 1701 210 1700 188 1699 225 1698 136 1697 143 1696 143 1695 119 227 39,3 0,92 0,35 0,11 0,00 0,02 0,52 1694 240 1693 120 246 39,4 1,10 0,34 0,07 0,00 0,01 0,54 1692 235 1691 121 187 40,9 0,95 0,50 0,06 0,00 0,03 0,56 1690 64 1689 188 1688 122 94 40,8 1,03 0,50 0,04 0,00 0,04 0,57 1687 265 1686 146 1685 167 1684 157 1683 179 Appendices 197 Appendices

Appendix 7.1: Photograph of gravity core PSS 94/3 and piston core PSS 94/4 taken from Lej da Segl in Septem- ber 1994. Note the lowermost 150 cm of non-laminated, homogenous sediment.

Lej da Segl (southwestern basin)

White label = Unit 3/4 boundary Altitude : 1797m a.s.l. Opening date : 14.03.1996 = Correlation between Water depth : 23m short-core and piston-core Core-length : 10.63m FSS 94/3 PSS 94/4 1 1 23 45 67 8910

10 10

20 20

30 30

40 40

50 50

60 60

70 70

80 3 80 4

90 90

100 100 Appendices 198 Appendices

Appendix 7.2: Median grainsize of core PCH 94/3 Late Glacial sediments (unit 4, 505-1159cm) compared to the mean and standard deviation of the grainsize distribution. The standard deviation is a measure of the sorting..

Grainsize Grainsize Grainsize Sed. median mean std.deviation depth [µm] [µm] [µm] [m] 3.0 4.0 5.0 6.0 3.0 5.0 7.0 9.0 11.0 1.0 3.0 5.0 7.0 9.0 11.0 5

4B 9

11 Appendices 199 Appendices

Appendix 7.3: Declination curve and intensity of the palaeomagnetic vector for core PCH 94/3.

Sed. Declination PCH 94/3 Intensity depth Lithologic Dates [°] [mA/m] [m] units 14C years BP 60 120 180 240 300 360 0 100 200 300 400 0 1 1 1200 b

2 2 2100 c

2740 3 d

4 4725 e 3 5000

5 9350

10000 Θ? 6 10065 4A 7 10370 11315 8

9 4B 11055 11000 I? 10330 (12400) (12800) I-j? 10 (13000) I-k? I-l? 4C (13500) I-m? 11 7260 17970 Acknowledgments 200 Acknowledgments Acknowledgments

First of all I would like to thank Prof. Dr. Helmut Weissert for his tolerance and patience. He educated me in positive thinking, which was sometimes not so easy. Besides the scientific advice I am grateful to Helmi because he always gave me the freedom to sort out the private problems I had during my thesis. Without his support and encouragement I would never have finished this thesis. Thank you very much for that Helmi.

The same is true for Dr. Frank Niessen. He represented the ‘northern element’ in my advisor team, which I, as a northern German in Switzerland, definitely needed from time to time. Discussing results with him always was a pleasure because of his enthusiasm and his high rate of producing new ideas. Every meeting with him took me a big step forward. Thank you very much Frank.

Important scientific and personal support I also received from Dr. Michael Sturm. Mike always made me re-think things by asking questions I did not think about any more, which improved my understanding of the topic. He gave me the opportunity to continue working in limnogeology. Thank you very much Mike.

Moving from Göttingen to Zürich sounds like taking a small step further south in Europe. When I came here I soon realized that the difference is much larger than is apparent from the map. I did not expect to meet people from all over the world in the geological institute of ETH, Zürich. I enjoyed meeting Jane, Trey, Andrea, Jan-Henk, Geoffrey, Alex, Jorijntje, Stephano, Bill, Flavio, Barbara, Daniel, Christina, Hermann, Catalina, all Markuses, Oli, Chrisogono, Steve, Graham, Mara, Piotr, all Ulis, Stefan, Jörg, Karsten, Abdel Asis, Achmet, Susan, Andrzej, Gerold, all Judithes, Steve, Iwan, Lisa, Gianreto, Maureen, Hilary, Ali, Milena, Andreas, Marco, Giulio, Mary, Wilfried, Diane, Johannes, Hanspeter, Dominique, Michael, Robert, Yvonne, Sylvia B., all Evas, Erika A. and Guy. Thank you all for being the way you are. I apologize to those, I forgot to mention and I’m sure that I forgot someone, I always do.

Another positive aspect of my thesis was the beautiful field area, the Engadine. All the people I met in the Engadine were always very friendly and helpful. In particular it was Toni Klucker who made sometimes rather spontaneous sampling campaigns possible, due to his unburocratic help. The communities of Silvaplauna and Segl are acknowledged because they never said no when I wanted to take samples. During several sampling campaigns, a lot of people offered helping hands. Field- work with Kurt, Peter, Robert, Daniel, Gerry, Andrzej, Trey, Bill and Caroline was very enjoyable. Thank you very much. I also thank Peter Baumann from the company LIMNEX for letting me use their water chemistry data.

I lived together with a Texan and an Italian above a disco and in Kreis 5 and enjoyed it very much. Do I have to say more ...... ? OK, sharing flat and office with Trey and Andrea opened new views on live and was a very important piece of education. In the office Jan-Henk was responsible for the Dutch perspective. Since I took time for my thesis I had the privilege to get to know a second generation of office mates. Alexandra, Jorijntje and Geoffrey are responsible for a pleasant office ‘climate’ making things easy for me, the office dinosaur. Acknowledgments 201 Acknowledgments

I learned a lot in discussions with Daniel Ariztegui, Christina Chondrogianni, Gerry Lemcke, Andreas Leemann, David Livingstone, Guy Lister, William Anderson, Trey Meckel, Andrea Artoni, Piotr Gawenda, Uli Wortmann, Martin Wessels, Willy Tinner, Prof. Katharina von Salis Perch-Nielsen, Prof. Judith McKenzie and PD Dr. Peter Hochuli. Thank you very much. I also thank Prof. Wilfried Haeberli and Dr. Daniel VonderMühll for introducing me to the geomorphology of the Upper Engadine and Dr. Felix Keller for giving me the opportunity to discuss my results with the local people in the Engadine at the Academia Egiadina. In the initial phase of the work Jürgen Schneider and Ulrike Wolf helped me to find my way into the thesis. Thank you for that.

For technical support and help during the fieldwork, I am grateful to Kurt Ghilardi, Robert Hofmann, Peter Holste, Andrzej Fischer and Anton Klucker. Caroline Stengel not only let me measure the carbon content in hundreds of virtually carbon free samples. Thanks to her helpfulness lab work never was a problem. I thank Wolfgang Bonn who taught me how to measure biogenic silica. Urs Gerber did all the photographic work in no time. I don’t know how he could do that but he certainly deserves a big ‘thank you very much’ for that. Dr. Stahel let me measure all of my samples on his XRD-machine. Miraculously he was always able to fix every problem that I caused by mis- usage of the software. Thank you very much. The magnetic measurements were possible thanks to the help of Prof. Heller and Thomas Forster. I also whish to thank Gianreto Manatschal, who translated the abstract into Rhätoromanisch.

Erika Gobet inherited my cores and will continue working on them for a pollenstratigraphy. Good look for that and thank you for the discussions we had and hopefully still will have in the future.

Considering my laziness in keeping in touch type-of-things I am very happy to have friends like Thomas Kulbe and Dorothee Barnikol-Schlamm. Sandra and Peter and Oma Lotte were there when help was needed or when the Engadine ‘called’ for a visit. Of course I would not have even made it to Zürich to start a thesis without the support of my parents. Zu wissen, daß, wenn nichts mehr geht und alles schief läuft, immer noch jemand da ist, zu dem man zurück kann, ist mehr als ein beruhigendes Gefühl. My Concu(Sa)bine accompanied me when I was running, pushed me when I stood still and caught me when I was stumbling. Thank you for maintaining our family life during times that were not always easy. The person that had the biggest impact on my thesis certainly was Lea, my daughter. She reminded me that the thesis is not the most important thing in the world.

Curriculum Vitae 203 Curriculum Vitae

Curriculum Vitae

Christian Ohlendorf

Born on April 26., 1966 in Buchholz i.d.N., Germany

Citizenship: German

Education:

1972 - 1974: Grundschule in Kampen, Germany. 1974 - 1976: Grundschule in Otter, Germany. 1976 - 1978: Realschule in Tostedt, Germany. 1978 - 1985: Gymnasium am Kattenberge in Buchholz i.d.N., Germany. Degree: Abitur. 1985 - 1986: Army service, Germany. 1987: Employment at Geoanalytik Labor und Consult GmbH, Hildesheim, Germany. 1987 - 1993: Studies and Diploma thesis in Geology at the Georg-August Universität, Göttingen, Germany. Supervisor: Prof. Dr. J. Schneider Degree: Diplom Geologe. 1993 - 1998: Doctoral student and research assistant at the Geological Institute of ETH, Zürich, Switzerland. Supervisors: Prof. Dr. H. Weissert, Dr. F. Niessen, Dr. M. Sturm.