Event History of the Santiago area (Chile): the sedimentological archive of Lago Lo Encañado

Thomas Pille

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Acknowledgements

Om te beginnen zou ik graag mijn promotor, Prof. Dr. Marc De Batist, bedanken voor dit interessante onderwerp en voor de tijd die u hierin gestoken hebt, voor de snelle verbeteringen en de veelbetekende tips en commentaren.

Ook Maarten, mijn begeleider, verdient een welgemeende bedanking. Terwijl je het dit jaar zelf heel druk had (een doctoraat dat afgemaakt moest worden, je eerste kindje, een nieuw onderzoeksproject) vond je toch altijd de tijd om mij bij te staan. Vooral je enthousiasme en je doorzettingsvermogen zullen me bijblijven. Voor ieder klein vraagje waarmee ik kwam waren we een paar uur bezig, en vertrok ik niet enkel met een antwoord, maar ook met een hoofd vol extra ideeën.

Philipp, obwohl ich ursprünglich nicht dein Thesis-Student war, hast du in der Abwesenheit von Maarten sehr viel Zeit in mein Project investiert. Was mir vor allem gefallen hat war das du bei einer Frage nicht einfach die Antwort gegeben hast, sondern auch versucht hast mir das Prinzip dahinter zu erklären. Vielen Dank für al deine Hilfe, und für eine schöne Reise nach Brest. Ich wünsche dir alles Gute bei deiner Doktorarbeit.

I would also like to thank the rest of the RCMG staff, especially to Thomas, Mario, Willem, Oscar, Stan, Katrien and Koen. You were always there to help me if needed, or for a chat during a coffee break. The RCMG is a fantastic working environment.

Mijn familie heeft er dit jaar (en eigenlijk doorheen mijn 5-jarige universitaire carrier) altijd voor mij gestaan. Bedankt voor de hulp, en bedankt om mij vanuit het buitenland de kans en het vertrouwen te geven om mijn eigen weg te zoeken.

Ook van mijn klasgenoten heb ik veel steun gekregen. In deze tijden van thesis-stress geldt dat vooral voor Stef, Jeroen en Tim. Ik vond de koffiepauzes altijd een zeer aangenaam moment om de sleur van de dag te doorbreken, en de late-night schrijfsessies in de S8 zouden niet gelukt zijn zonder de steun die we van elkaar kregen.

Ten slotte mag ik zeker mijn vriendin Roosje niet vergeten. Jij stond altijd achter mij, luisterde naar mijn gezaag en pepte mij op als ik het even niet meer zag zitten. Bedankt Roosje, jij bent de beste!

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Nederlandstalige samenvatting

Tijdens twee veldcampagnes in 2011 en 2012 werden in totaal negen sedimentkernen genomen in Lago Lo Encañado, een klein meertje in de Chileense Andes op ongeveer 50 km ten oosten van de hoofdstad Santiago. Van die negen kernen werden er acht geopend en bestudeerd, met als bedoeling het vinden van sporen van aardbevingen. Chili, en bij uitbreiding de hele Zuid-Amerikaanse westkust, is een tectonisch zeer actief gebied. Deze kust bevindt zich langs een subductiezone waar de oceanische Nazca plaat (en in het zuiden van Chili de Antarctische plaat) onder de continentale Zuid-Amerikaanse plaat duikt. Deze platenbeweging kan zeer zware subductie- (of megathrust) aardbevingen produceren. In Chili komen subductie-aardbevingen met een moment magnitude (Mw) > 8.5 gemiddeld een paar keer per eeuw voor (Udías et al., 2012). Deze aardbevingen hebben in veel gevallen een hoge dodentol tot gevolg, omdat ze vaak ook een tsunami opwekken. Bekende voorbeelden zijn onder andere de Maule aardbeving in 2010 (Mw 8.8; Wang et al., 2012) en de Valdivia aardbeving in 1960 (Mw 9.5; Kanamori, 1977). Deze laatste is wereldwijd de krachtigste aardbeving ooit waargenomen. In centraal en zuid Chili kan men de subductiezone verdelen in drie segmenten, met van noord naar zuid het Valparaíso segment, het Concepción segment en het Valdivia segment. Naast subductie-aardbevingen brengt het subductieproces nog andere gevaren met zich mee, waaronder intraplaat-aardbevingen en vulkaanuitbarstingen. Intraplaat-aardbevingen worden veroorzaakt door spanning op de bovenliggende continentale korst (ondiepe aardbevingen) of op de onderduikende oceanische korst (diepe aardbevingen). Deze aardbevingen zijn over het algemeen gekenmerkt door een veel lagere magnitude, maar kunnen lokaal minstens even schadelijk zijn. De dodelijkste aardbeving in de geschiedenis van Chili was een intraplaat-aardbeving in 1939 in het stadje Chillán (Mw 8.3). Deze aardbeving kostte het leven aan 30.000 mensen. Het hoge aantal krachtige aardbevingen zorgt ervoor dat Chili wereldwijd een van de landen is met het hoogst aantal doden per aantal inwoners als het gevolg van geohazards (Lomnitz, 1970a).

Om een beter zicht te krijgen op de kans op en de kracht van toekomstige aardbevingen is het van belang om, naast real-time observaties, de tektonische geschiedenis van dit land en deze regio te kennen. Deze is echter onvolledig gekend en dat heeft twee belangrijke redenen. Ten eerste bestaan er maar historische documenten vanaf de aankomst van de Spaanse veroveraars in 1541. De oorspronkelijke bewoners, de Mapuche, hadden namelijk geen schrift. Daarnaast zijn grote delen van Chili zeer dun bevolkt of helemaal onbewoond, wat het aantal waarnemingen sterk limiteert. Enkel in de grote steden vindt men grote bevolkingsdichtheden. Zo is de regio rond Santiago (Región Metropolitana de Santiago) met meer dan 6 miljoen inwoners goed voor bijna 40% van de bevolking van het land.

Om meer te weten te komen over de tektonische geschiedenis van dit land is het nodig om andere bronnen dan historische waarnemingen te raadplegen. Zo zijn er bijvoorbeeld

5 onderzoekers die dit proberen te doen aan hand van tsunami-afzettingen (Cisternas et al., 2005). Deze methode blijkt echter minder gevoelig te zijn als men dit vergelijkt met gekende aardbevingen, aangezien niet iedere aardbeving een tsunami veroorzaakt. Een archief dat wel nauwkeurig en gevoelig is, en al in verschillende aardbevingsgevoelige gebieden gebruikt is (Howarth et al., 2012 in Nieuw Zeeland; Beck, 2009 in de Franse alpen), zijn meersedimenten. Het doel van deze thesis is voornamelijk het vinden van aardbevingen in dergelijke meersedimenten en het testen van de gevoeligheid van dit specifieke meer op aardbevingen.

Het onderzoeksgebied is een klein meertje in de Andes van Centraal Chili. Deze locatie is zeer interessant omdat het op 50 km van het grootste bevolkingscentrum van Chili ligt. Lago Lo Encañado is een proglaciaal meer op een hoogte van 2500 boven de zeespiegel in de Maipo vallei. Er bevindt zich nog steeds een gletsjer in het noordelijkste uiteinde van het drainagegebied. Bijzonder aan dit meer is dat het drainagegebied (30 km²) veel groter is dan het meer zelf (0.5 km²). Het drainagebekken kent bovendien een zeer groot hoogteverschil. Zo is het hoogste punt gelegen op meer dan 4100 m boven zeespiegel. Deze factoren zorgen ervoor dat Lago Lo Encañado een grote en bovendien zeer klastische sedimenttoevoer kent. Het grootste deel van het sediment wordt aangevoerd door een riviertje, de Río lo Encañado, dat door de glaciale vallei stroomt, ten noorden van het meer. De gemiddelde sedimentatiesnelheid bedraagt 1.75 mm/jaar (Salvetti, 2006).

Om het meer te bestuderen werden acht boorkernen geopend, beschreven, gemeten en bestudeerd. Met behulp van een Multi-Sensor Core logger (MSCL) (UGent) werd de magnetische susceptibiliteit gemeten (Bartington puntsensor), de densiteit bepaald met behulp van gammastralen en spectrofotometrie gedaan. Het sediment werd vervolgens bestudeerd aan hand van kleurbewerkte fotos en er werden smeerplaatjes gemaakt voor een microscopische studie. Om interessante laagjes te kunnen dateren werden de kernen gecorreleerd met een kern (kern 1305) uit een andere studie (Salvetti, 2006). Deze kern is gedateerd met behulp van 210Pb voor het bovenste deel (tot 1906 AD) en met 14C voor het onderste deel. Omdat de 14C dateringen onrealistisch grote ouderdommen opleveren werden deze verworpen, en werd een ouderdomsmodel ontwikkeld dat gebaseerd is op een extrapolatie van de sedimentatiesnelheden bepaald in de bovenste sedimenten. Het meer werd in kaart gebracht door middel van een sidescan-sonar survey. Hiermee werd een reflectiviteitskaart en een bathymetrische kaart gemaakt. Op beide kaarten zijn een aantal structuren zichtbaar. Onder andere kan in het noorden van het meer een delta onderscheiden worden, waarop kanaaltjes gevormd zijn waarlangs het sediment van de rivier het meer binnenkomt. Ook zijn de steile hellingen van de oostelijke en westelijke randen zichtbaar, net als een aantal puinwaaiers die op verschillende plaatsen het meer in gaan. Deze puinwaaiers bevatten blokken van meer als 1 m diameter, die op de reflectiviteitskaart zichtbaar zijn.

In de kernen werden aanwijzingen voor aardbevingen gevonden. Deze bestaan uit afzettingen van laagjes met een grove, zandige, basis, die naar boven toe fijner worden. Deze laagjes worden vaak bedekt met een wit laagje van variabele dikte. De totale dikte van zulke laagjes

6 kan variëren van enkele mm tot meer dan 20 cm. Deze laagjes kunnen tussen de verschillende boorkernen gecorreleerd worden. In het algemeen geldt dat deze laagjes een grotere korrelgrootte hebben naarmate ze zich dichter bij de delta bevinden. Zo heeft de eerste kern (ENC01) een basis die bestaat uit grind. Daar zijn deze lagen erosief, wat ervoor zorgt dat de stratigrafie niet compleet is. Ook de dikte van de laagjes is groter met een kleinere afstand van de delta, of van de steile rotswanden aan de oostelijke en westelijke rand van het meer. Onder een turbidiet is het sediment soms verstoord. Dat is voornamelijk het geval in de kernen het dichtst bij de randen van het meer. Deze verstoring kan gaan van een plooiing van het sediment tot een volledige homogenisatie van het sediment. Deze verstoringen zijn geïnterpreteerd als afzettingen van een afglijding of massastroom (de zogenaamde Mass- Transport Deposits of MTDs). De grofkorrelige afzettingen werden geïnterpreteerd als turbidieten, die in dezelfde beweging afgezet zijn als de MTDs. Microscopisch onderzoek toont aan dat deze turbidieten bestaan uit klastisch, terrigeen materiaal. Dit type van turbidieten heeft niet met zekerheid een seismische oorsprong (Van Daele, 2013). Een delta kan ook instorten onder zijn eigen gewicht, en de steile oostelijke en westelijke bergflanken kunnen ook zonder externe oorzaak in het meer afglijden. Ook een stortvloed kan grofkorrelig materiaal in het meer transporteren. Er zijn echter enkele aanwijzingen die in het voordeel van een seismische oorsprong spreken. Zo zijn de MTD’s onder de turbidieten het gevolg van een afglijding en zijn dus niet verklaarbaar door een stortvloed. Daarnaast zijn er in enkele turbidieten meerdere cycli (grofkorrelige basis, fijnkorreliger naar de top) te zien, onder andere in ENC08. Dat wijst op verschillende afglijdingen die vlak na elkaar gebeurd zijn, of op verschillende afglijdingen die op hetzelfde moment op verschillende plaatsen binnen het meer gebeurd zijn. In beide gevallen is een aardbeving de enig mogelijke verklaring.

Omdat het sediment in de kernen in de buurt van de kustlijn sterk vervormd is, en omdat de turbidieten daar vaak erosief zijn, werd de ‘event’-stratigrafie vooral gebaseerd op de kernen in het centrale deel van het meer, met name ENC02 en ENC03. Op die manier werden in totaal 13 dergelijke laagjes met een grofkorrelige basis aangetroffen, waarvan er acht konden gelinkt worden aan een historische aardbeving. Voor elke aardbeving werd een seismische intensiteit berekend. Dit werd gedaan met behulp van formules in Bakun & Wentworth (1997) voor lokale aardbevingen, en in Barrientos (1980) voor subductie-aardbevingen. Deze formules geven een theoretische seismische intensiteit op basis van de moment magnitude Mw (lokale aardbevingen) of de oppervlaktegolf-magnitude Ms (subductie-aardbevingen) en de afstand tot het epicenter (lokale aardbeving) of tot aan de breukzone (subductie-aardbeving). De waarden zijn uitgedrukt op de Modified Mercalli Intensity (MMI) schaal. Hiermee wordt de schade die een aardbeving op een bepaalde plaats kan veroorzaken beschreven. Er bestaat geen algemeen aanvaarde drempelwaarde in seismische intensiteit waarboven een seismisch- geïnduceerde afzetting gevormd wordt in een meer. Enkele studies geven echter een empirische drempelwaarde van VI (Van Daele, 2013) en VI-VII (Monecke et al., 2004). Éen van de doeleinden van deze MSc thesis is om te achterhalen hoe goed deze drempelwaarde overeenstemt met de werkelijkheid in Lago Lo Encañado. In de rest van dit hoofdstuk zal de

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MMI gegeven worden in Arabische cijfers in plaats van de gebruikelijke Romeinse cijfers, omdat de gebruikte formules vaak komma-getallen opleveren.

De bovenste turbidiet (event A) komt enkel voor in de kernen die genomen zijn tijdens de tweede veldstage. Dat geeft ons dus een tijdskader waarbinnen de aardbeving gebeurd is (tussen januari 2011 en februari 2012). De enige aardbeving die hiervoor in aanmerking komt is een Mw 4.0 lokale aardbeving, op 15 km ten zuiden van het meer. De berekende MMI (5) is lager dan de verwachte drempelwaarde, en er zou dus geen turbidiet mogen gevormd worden. Mogelijk is dit te verklaren door een uitzonderlijk lage waterspiegel op het moment van de aardbeving. Hierdoor wordt een instabiele situatie gecreëerd (Moernaut et al., 2010), en in deze setting is de turbidiet bij een relatief lage intensiteit wel te verklaren.

Event B is de grootste turbidiet in de kernen, met een maximale dikte van 23 cm (ENC01). Onder andere in ENC05 en in ENC09 is deze turbidiet afgezet samen met een MTD. Deze laag is echter niet zichtbaar in kern 1305 (uit Salvetti, 2006). Dat wil zeggen dat deze turibidiet gevormd is tussen 2011 en 2005. De meest voor de hand liggende oorzaak is de krachtige (Mw 8.8) Maule aardbeving in 2010. De MMI werd berekend op 7.1.

Op basis van de al beschreven 210Pb datering in Salvetti (2006) konden events C, D en E gedateerd worden. Event C wordt toegeschreven aan een subductie-aardbeving op het

Valparaíso segment in 1985 (Mw 7.8). De seismische intensiteit voor deze aardbeving bedraagt 6.2. Event D ligt net onder een piek in radioactieve neerslag. Deze is het gevolg van atmosferische atoomtesten tijdens de Koude Oorlog, en bereikte een hoogtepunt in 1962. In de zuidelijke hemisfeer valt de piek in neerslag van radioactieve deeltjes in 1964. Aangezien event D dieper ligt dan deze piek, zou men verwachten dat deze turbidiet gevormd is door de

Valdivia aardbeving in 1960. Met een Mw van 9.5 is dat wereldwijd de zwaarste aardbeving van de 20ste eeuw. Het Valdivia-segment ligt echter zo ver naar het zuiden dat de seismische intensiteit voor Lago Lo Encañado maar 4.9 is. Een betere verklaring voor event D is een lokale aardbeving in 1958. Deze aardbeving in de Maipo vallei had maar een Mw van 6.3, maar met een epicenter op 18 km van het meer levert dit een seismische intensiteit van 7.3 op. De volgende turbidiet, gelinkt aan event E, heeft een gedateerde ouderdom van ongeveer 100 jaar voor 2005 (toen kern 1305 gedateerd werd). Dit valt ongeveer samen met een aardbeving

(Mw 8.6) uit 1906 op het Valparaíso segment.

Onder deze turbidiet is er geen datering meer beschikbaar, omdat de hoge 14C ouderdommen verworpen werden. Daarom werd een ouderdomsmodel ontworpen op basis van sedimentatiesnelheden. Omdat de kleur en dikte van de achtergrondlaminaties plots verandert onder event I werd het model zo aangepast dat de sedimentatiesnelheid 0.167 cm/jaar bedraagt onder dit event, en 0.294 cm/jaar erboven. Op die manier konden H, I, K en M gelinkt worden aan aardbevingen in respectievelijk 1835, 1822, 1751 en 1730. Al deze aardbevingen hebben een MMI hoger dan de vooropgestelde drempel van 6.

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Er zijn nog vier events die niet gelinkt konden worden aan gekende aardbevingen. Twee hiervan, events F en G, zijn afgezet vlak na elkaar waardoor ze een ongeveer gelijke ouderdom hebben. Deze hebben een gemodelleerde ouderdom van respectievelijk 1892 (voor F en G), 1800 (J) en 1746 (L). Er zijn een aantal mogelijkheden om dit te verklaren. Het zou kunnen gaan om niet gekende, lokale aardbevingen in de Maipo vallei. De vallei is zeer dunbevolkt, en het is dus zeker mogelijk dat er aardbevingen niet waargenomen zijn. Een tweede optie is dat er, vergelijkbaar met de lage waterspiegel in 2012, omstandigheden opgetreden zijn waardoor de drempelwaarde waarboven een turbidiet gevormd wordt tijdelijk verlaagd was. Zowel voor event F+G en event J was er een periode van 50 jaar of langer zonder turbidieten. Het kan dus zijn dat er in de tussentijd zoveel sediment afgezet is op de delta dat die al onstabiel werd, en een kleine aardschok voldoende was om een afglijding te veroorzaken. Een laatste optie is dat deze laagjes verkeerdelijk als turbidieten met een aardbeving als oorzaak geïnterpreteerd zijn.

Van alle aardbevingen werd een seismische intensiteit berekend, en op één na alle historische aardbevingen met een geschatte intensiteit van VI of hoger zijn geregistreerd in dit meer. Enkel van een lokale aardbeving in 1850 zijn geen sporen te vinden in het sedimentarchief. Nochtans zou dit event de hoogste seismische intensiteit moeten hebben. Waarschijnlijk ligt de oorzaak hiervan ook in de afgelegenheid van deze vallei, en zijn de magnitude en/of het epicenter verkeerd ingeschat. Afgezien van deze onzekerheid kan geconcludeerd worden dat Lago Lo Encañado een gevoelig archief is om aardbevingen te registreren. Het is daarom een zeer geschikt meer om verder onderzoek op te doen. Door diepere kernen te nemen en te dateren zou het mogelijk moeten zijn om bewijzen van aardbevingen te vinden van voor de tijd van de Spaanse kolonisatie.

Een vergelijking met seismische ‘event’-stratigrafieën in andere meren in de omgeving leert dat deze meren een iets hogere intensiteitsdrempel hebben dan Lago Lo Encañado, maar dat deze nog binnen de schatting van Monecke et al. (2004) vallen (VI-VVI). De meren zijn Laguna del Inca, Laguna Negra, Laguna el Ocho en Laguna Aculeo. De data hiervoor komen uit enkele paleoklimatologische studies (von Gunten, 2009a; 2009b; 2009c; Vandenberghe, 2012). Door ‘event’-stratigrafieën uit verschillende meren met elkaar te vergelijken kan men meer te weten komen over het type aardbeving (subductie-aardbeving of lokale aardbeving). Op een lang noord-zuid traject zou men zelf het segment van een onbekende subductie-aardbeving kunnen achterhalen. De meest interessante onderwerpen voor een verdere studie liggen in het vinden en beschrijven van aardbevingen van voor de aankomst van de Spanjaarden. Dit vergt echter diepe kernen uit verschillende meren, die elk absoluut kunnen gedateerd worden. Deze studie heeft aangetoond dat het sedimentarchief van Lago Lo Encañado voor een dergelijk onderzoek geschikt is.

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Table of contents

Acknowledgements

Nederlandstalige samenvatting

List of Figures

1. Introduction...... 15

2. Lake sediments...... 17

3. Seismology...... 21 3.1 “Hazard” vs. “Risk”……………………………………...... 21 3.2 What is a fault segment...... 21 3.3 Magnitude and intensity scales...... 23 3.3.1. Magnitude scales...... 23 3.3.2. Intensity scales...... 24

4. Setting...... 27 4.1 Geographical setting...... 27 4.2 Geological setting...... 30 4.2.1. Tectonic setting...... 30 4.2.2. Local geology...... 34 4.3. Historical setting...... 36

5. The historical archive of major earthquakes in Central Chile...... 39 5.1 Subduction earthquakes...... 39 5.1.1. Valparaíso segment...... 39 5.1.2. Concepción segment...... 41 5.1.3. Valdivia segment...... 43 5.2 Intraplate earthquakes...... 46 5.3 Earthquakes with an unknown source...... 47

6. Previous studies on Lago Lo Encañado and nearby lakes...... 49 6.1 Lago Lo Encañado...... 49 6.2 Surrounding lakes...... 50

7. Methods...... 53 7.1 Coring...... 53 7.2 Geoacoustics...... 54 7.2.1. Sidescan sonar...... 54 7.2.2. Sonarscope...... 55 7.3 Core opening...... 56 7.4 Non-destructive core analysis...... 57 7.4.1. MSCL analysis...... 57 7.4.1.1. Magnetic susceptibility...... 58 7.4.1.2. Colour spectrophotometry...... 58 7.4.1.3. Gamma-ray attenuation density...... 59 7.5 Destructive core analysis...... 59

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7.5.1. Smear slides...... 59

8. Results...... 61 8.1 Lake floor morphology: observations and interpretations...... 61 8.2 Sediment cores: observations, analysis and interpretation...... 69 8.2.1. Background sediment...... 70 8.2.2. Fining-upwards layers...... 70 8.2.3. Blue layers...... 72 8.2.4. Red layers...... 72 8.2.5. White layers...... 73 8.3 Core descriptions...... 73

9. Interpretation and discussion...... 79 9.1 Interpretation of the different layers...... 79 9.2 Correlation of the turbidites...... 84 9.3 The intensity of an earthquake...... 87 9.4 Historical earthquakes and their deposits...... 90 9.4.1. 20th and 21st century earthquakes...... 90 9.4.2. Comments on the 20th and 21st century earthquakes...... 94 9.4.3. Pre-1900 earthquakes...... 94 9.4.4. Comments on the pre-1900 earthquakes...... 97 9.5 Comparing Lago Lo Encañado to other lakes in the proximity...... 97 9.6 Volcanic activity...... 103

10. Conclusion and outlook...... 105 10.1 Conclusion...... 105 10.2 Outlook...... 106

11. References...... 109

12. Attachments...... 119

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List of figures

2.1 The four types of density flows...... 18 2.2 The Bouma sequence...... 19 3.1 Relation between moment magnitude and several other magnitude scales...... 24 4.1 Map of South America and location of the lake...... 28 4.2 Photo of Lago Lo Encañado...... 29 4.3 Satellite image of Lago Lo Encañado...... 30 4.4 Terrain map of South Chile...... 32 4.5 An overview of the major earthquake segments along the western South-American coast 33 4.6 Geological map of Central Chile...... 35 4.7 Elevation map of the Santiago region with the position of the closest volcanoes...... 36 5.1 Map of Central Chile with locations of cities and rivers along the Valparaíso segment...... 40 5.2 Map of Central Chile with locations of cities along the Concepción segment...... 42 5.3 Map of South-Central Chile with locations of important geographical features on the Valdivia segment...... 45 5.4 Map of the Maipo Valley...... 47 6.1 Core 1305, core log and age-depth model...... 50 6.2 Map of the Santiago area with the location of other lakes mentioned in this study...... 51 7.1 Satellite image of Lago Lo Encañado showing the core locations...... 54 7.2 Sidescan-sonar survey lines with heading...... 56 7.3 Sketch of a Geotek Multi-Sensor Core Logger...... 58 8.1 Reflectivity map of Lago Lo Encañado...... 62 8.2 Bathymetry map of Lago Lo Encañado...... 63 8.3a Details of the reflectivity map: central structure...... 65 8.3b Details of the reflectivity map: colluvial fan...... 65 8.3c Details of the reflectivity map: delta channels...... 66 8.3d Details of the reflectivity map: high reflectivity line...... 67 8.4 View from the eastern shore on Lago Lo Encañado...... 68 8.5 Interpretation of the sidescan-sonar map...... 69 9.1 A comparison between the Bouma sequence and turbidites in this thesis...... 79 9.2 A comparison between cores from Laguna Negra and Lago Lo Encañado...... 82 9.3 A correlation of the turbidites in the cores on the southern line...... 85 9.4 A correlation of the turbidites in the cores on the southern line...... 87 9.5 A comparison between ENC03, ENC02 and 1305...... 91 9.6 Seismic intensity map of the 1958 Maipo Valley earthquake...... 93 9.7 Age-depth model...... 95 12.1 Core log of ENC01...... 120 12.2 Core log of ENC02...... 121 12.3 Core log of ENC03...... 122 12.4 Core log of ENC05...... 123 12.5 Core log of ENC06...... 124 12.6 Core log of ENC07...... 125 12.7 Core log of ENC08...... 126 12.8 Core log of ENC09...... 127 12.9 Legend to the core logs...... 128

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List of Tables

7.1 Core lengths and locations and year of coring...... 53 8.1 A recapitulation of the fining-upwards layers in ENC01 and ENC02...... 74 8.2 A recapitulation of the fining-upwards layers in ENC03 and ENC05...... 76 8.3 A recapitulation of the fining-upwards layers in ENC06 and ENC07...... 77 8.4 A recapitulation of the fining-upwards layers in ENC08 and ENC09...... 78 9.1 Summary of all known historic earthquakes in Central Chile...... 89 9.2 Calculated MMI for Lago Lo Encañado and surrounding lakes...... 101-102 9.3 List of all known eruptions of three nearby volcanoes...... 103

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1. Introduction

Earthquakes are a very common geohazard in Chile. In fact, the country has one of the highest casualties per capita ratio caused by earthquakes and earthquake-induced tsunamis in the last century (Lomnitz, 1970a). Earthquakes occur on a daily basis, and most of them are harmless because they are of low magnitude and/or occur at great depth. There were over 150 earthquakes with a Mw (moment magnitude) above 4.0 along the South American west coast in 2012 (USGS, 2013a). Strong earthquakes with a Mw > 8.5 occur at least a few times per century (Udías et al., 2012). These earthquakes cause vast damage and loss of life, not only in Chile, but through tsunamis along the coasts of Pacific Ocean as well. The most recent significant earthquake was the Maule earthquake on 27 February 2010 (Mw 8.8). This earthquake ranks in the top-10 of the world’s largest recorded earthquakes, although the event stood in the shadow of the smaller, but far deadlier Haiti earthquake on 12 January

2010. On 22 May 1960, Chile was struck by the strongest earthquake ever recorded. This Mw 9.5 Valdivia earthquake produced a tsunami that not only cost life and destroyed coastal towns in Chile (Sievers, 2000), but also hit Hawaii (10 m wave, Lander & Lockridge, 1989) and Japan (6 m wave, Watanabe, 1998; in Cisternas et al., 2005). On 20 February 1835, Charles Darwin experienced an earthquake while in Chile during his voyage of the Beagle. He describes the earthquake and its consequences in Valdivia and Concepción in an objective yet vivid way (Darwin, 1845).

At present, earthquakes are still unpredictable. The concepts of plate tectonics, continental drift and seafloor spreading have helped to improve our understanding of the large-scale movements of the lithosphere, and areas prone to earthquakes are well known. For example, the coast south of Santiago along which the 2010 Maule earthquake ruptured was identified as a seismic gap in the 1990’s (Campos et al., 2002; Madariaga et al., 2010). This case is an example of how scientific research has yielded some positive results. Despite all these efforts, many strong earthquakes still occur unexpectedly. While writing this thesis, this was made clear again by an earthquake in the border region between Iran and Pakistan on 16 April 2013. Examples of unpredicted earthquakes in a similar subduction setting as in Chile are the 2011 Tohuko-Oki and the 2004 Sumatra earthquake (Meltzner et al, 2012). Many theories to predict earthquakes have been tested and abandoned (Hough, 2009; in Bilham, 2010). It is possible to identify where catastrophic events will occur, but it remains impossible to predict the time of the event (Madariaga et al., 2010). However, it is a global ongoing scientific effort to assess geo-hazards, like earthquakes, and their risks. These studies are based on two pillars. The first pillar comprises real-time measurements like plate movement, locking between two plates, uplift or subsidence. Moreno et al. (2011), for example, evaluate the dangers by monitoring plate motion and state of locking between the Nazca and South-American Plate in South-Central Chile with velocities derived by Global Positioning System (GPS) and kinematic

15 finite-element (FE-) models. The second pillar relies on studying evidence from the past. By investigating both historical and geological evidence, researchers are trying to find earthquake recurrence intervals. In Chile, Udías et al. (2012) and Lomnitz (2004) have scoured Spanish archives for the most important historical earthquakes. On the other hand, Cisternas et al. (2005) are also reconstructing earthquake history in South-Central Chile, but do that with geological evidence, in this case tsunami deposits and other estuarine evidence. The recurrence interval (nearly 300 years) found in tsunami deposits from the estuary of Rio Maullín is almost twice as high as the interval in the historical record (128 years). This means that not every event was recorded in the studied sedimentary archive.

Sensitivity is of great importance in paleoseismological studies. If an archive is known to record every earthquake above a certain threshold, it provides a solid basis for recurrence calculations. Lake sediments have proven to be a useful archive because they are sensitive recorders of earthquake activity and can contain a continuous and temporally well-resolved record, with a high preservation potential. Turbidites can be used for paleoseismological studies (e.g. Moernaut et al., 2007). Earthquakes have been inferred from lake sediments all over the world, for example along the Alpine fault in New Zealand (Howarth et al., 2012), in the north-western Alps (Beck, 2009) or along the North Anatolian Fault (Hubert-Ferrari et al., 2012). In this MSc thesis we will work on cores from Lago Lo Encañado in the Andes east of the metropolitan area of Chile. Our aim is to find and describe evidence of seismic events that have occurred in the last hundreds of years by studying sedimentary records on different locations in the lake. These events can be linked to earthquakes found in historical archives. This comparison will show the sensitivity of this lacustrine archive towards earthquakes, both in local seismic intensity and in type of earthquake.

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2. Lake sediments

Lacustrine sediments can be composed of autochthonous or of allochthonous sediments. Autochthonous sediments are formed inside the lake and consist mainly of organic matter and biogenic silica (e.g. algae and diatom tests). Allochthonous sediments originate outside the lake and are transported into the lake by ice, rivers, wind or gravitation. When this material enters the lake through fluvial transport, it can cause density flows.

Density flows and their deposits are discussed in more detail in this chapter. Subaqueous density flows are important mechanisms of sediment transport on slopes in both (deep-sea) marine and lacustrine settings. Density flows are complex and highly variable, and so are their deposits. It is therefore not suprising that there have been many definitions and classifications in literature. One classification divides density flows into cohesive (e.g. debris or mud flows) and non-cohesive density flows (Mulder & Alexander, 2001). Non-cohesive density flows are further differentiated based on the difference between the density of the flow ρf and the density of the water ρw. Three end-members can be defined (Fig. 2.1). Hypopycnal flows (ρf <

ρw) are known as overflows. They are important in river estuaries where the sediment load is dispersed as bouyant plumes. Homopycnal flows (ρf = ρw) have the same density as the water.

Hyperpycnal flows (ρf > ρw) or underflows are responsible for the transport of large volumes of sediment. In stratified water bodies with water layers of different densities (ρw1 and ρw2) a fourth end-member can be identified, a mesopycnal flow (when ρw1 < ρf < ρw2). This is a simplified classification as the type of flow can change along its course.

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Figure 2.1 The four types of density flows. ρf is the density of the flow, ρw is the density of the water (Mulder & Alexander, 2001).

The term turbidite is a loosely defined term which is commonly used to describe the deposits produced by a specific type of density flows (Shanmugam, 2000). A true turbidity current has a concentration of <9 volume% of particles and is supported by fluid turbulence (Bagnold, 1962; in Mulder & Alexander, 2001). The classic model for describing turbidites is the Bouma Sequence (Bouma, 1962; in Shanmugam, 2000) (Fig 2.2). A typical turbidite is divided into five units (Ta, Tb, Tc, Td and Te; Bouma, 1962). The base of the turbidite (Ta) may be erosive and is composed of massive to normally graded sandstones, possibly with pebbles at the base. Unit

Tb is laminated and contains fine- to medium-grained sandstones. Unit Tc contains ripple or

18 wavy laminations. The unit is composed of fine sand. Unit Td is a laminated siltstone. The top layer Te is a mudstone that is homogenous or laminated. This model was later improved for fine-grained turbidites (Stow & Shanmugam, 1980) and for coarse-grained turbidites (Lowe, 1982; in Shanmugam, 2000) (Fig. 2.2). These coarse-grained turbidites (also known as deposits of high-density turbidity currents or Lowe sequences) are not deposited by true turbidity currents (Mulder & Alexander, 2001), but a more detailed description of the physical flow properties would go beyond the scope of this thesis. This work deals with the deposits and the triggers of density flows. Therefore all non-cohesive density flow deposits will be described as turbidites.

Figure 2.2 The Bouma sequence (Bouma, 1962) in the middle, together with a more detailed description for coarse-grained turbidites (Lowe, 1982; left) and for fine-grained turbidites (Stow & Shanmugam, 1980; right). From Shanmugam (2000). There are several possible processes that can generate density flows. An earthquake can potentially trigger the collapse of a slope or delta in a lacustrine basin. The deposits formed by earthquake-triggered slides are grouped under the collective noun “seismites”. Delta collapses are, however, not per definition earthquake induced, as they can also occur in settings where strong earthquakes are not common. This is demonstrated by a large collapse in Lake Brienz in the spring of 1996. The delta collapsed under its own weight (Girardclos et al., 2007). However, in areas where the recurrence of earthquakes is shorter than the time needed for a delta to accumulate enough sediment to become unstable, turbidites are a good proxy for seismic activity (Van Daele, 2013).

The classification in Van Daele (2013) will be used to define the assumed seismites. This classification divides seismites into three categories, mass-transport deposits (MTD’s), in-situ

19 deformations and turbidites. Turbidites are further divided into lacustrine turbidites type 1 and type 2 (resp. LT1 and LT2) and homogenites.

MTD’s are sediment packages initially deposited on a slope that slide downslope as a whole. They often have an erosional base and the original lamination can be preserved or deformed. They can sometimes be recognized by a repetition of layers. MTD’s are the result of cohesive sediment flows according to the definition of Mulder & Alexander (2001).

In-situ deformations are changes to the original lamination (e.g. folding or inclination), without movement of the sediment package. These deformations generally occur in areas of the lake where a slope is absent. As there is no gravitational force, the trigger is almost certainly an earthquake.

Lacustrine turbidites type 1 (LT1) are mass-wasting deposits from gentle slopes. They contain hemipelagic sediments that do not differ from the background sediment. LT1 turbidites almost certainly have a seismic trigger. A typical LT1 turbidite consists mainly

of a Ta unit at the base and a Te unit towards the top according to the Bouma sequence.

Lacustrine turbidites type 2 (LT2) have a more terrigenous content. They are formed as a result of delta collapses or onshore landslides that slide into the lake. Deltas are known to be more unstable and therefore more vulnerable to collapses than gentle slopes with hemipelagic sediments. This is mainly due to a higher sedimentation rate (Strasser et al., 2013). LT2s are often coarse grained, dark in colour and have high MS values. LT2 turbidites do not per definition have a seismic trigger. Homogenous LT2

turbidites can consist of Ta (or even S3 units in the Lowe sequence close to the delta)

and Te units. Well-graded LT2’s can further exist of Tb, Tc and Td units. In chapter 9.1. an LT2 turbidite is compared to the Bouma and Lowe sequence.

A homogenite is a turbidite formed by mass wasting. It consists of a coarse layer at the base of the deposit, followed by a thick unit of homogenous mud. This is often covered by a clay or fine-silt cap. A homogenite has a seismic trigger, and can be formed within the lake or by a landslide surging into the lake.

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3. Seismology

3.1 “hazard” versus “risk” Hazard and risk are often falsely used as synonyms, which is why a clear definition is necessary (Tinti, 2012):

A danger is a phenomenon that could lead to loss of human life or property. An example of a danger is an earthquake.

A hazard is the probability that this danger will occur in an area within a given period of time.

The vulnerability measures the degree of damage to structures (for example a house or a bridge) or people as a result of a danger.

The risk is an expression of the expected impact of this danger occurring in a given area within a given period of time.

The previously mentioned earthquakes in Haiti (Mw 7.0) and Maule (Mw 8.8) in January and February of 2010, respectively, are a good example to clarify these definitions. As both countries are situated near a fault zone, they are both in danger of experiencing an earthquake. As there are a much higher number of large earthquakes in the subduction setting of Chile, the earthquake hazard in Chile is much higher than in Haiti. However the population and the poorly built houses in Haiti are more vulnerable to earthquakes. The earthquake hit the densely populated capital of Port-au-Prince, where tens of thousands died in collapsing houses (Kijewski-Correa & Taflanidis, 2012). Chile on the other hand is sparsely populated and houses are purpose-built to resist earthquakes (Madariaga et al., 2010). Although the hazard of an earthquake is higher in Chile, the risk is higher in Haiti because the impact of the earthquake hitting a vulnerable city like Port-au-Prince is a higher than an earthquake in Chile.

3.2 What is a fault segment? Long active fault zones only rupture over a certain distance during a single earthquake. One hypothesis explains this by dividing faults into consistent fault segments. This means that earthquakes are always restricted to the segment where the rupture began. As a consequence, the maximum magnitude of an earthquake on a single segment is limited, and can be estimated even without knowledge of historical earthquakes on the same segment, according to this hypothesis. A part of a fault that has demonstrably ruptured to the surface at least twice is an earthquake segment by the definition of dePolo et al. (1989; In McCalpin, 1996). The evidence has to come from either documented historic surface ruptures or from multiple prehistoric ruptures that were limited to the segment. Earthquake rupture segments

21 or rupture segments are occasionally used as synonyms. A more general term is fault segment, which describes a part of the fault that can be distinguished from other parts of the fault in terms of earthquake behavior. A segment boundary is defined in Wheeler (1989) as “a portion of a fault where at least two (preferably successive) rupture zones have ended”. A segment boundary is seen as being persistent if a single rupture zone spanned a fault segment adjacent to the boundary and the rupture did not cross the boundary. The concept of fault segmentation is still under debate, because some case studies suggest that rupturing is in reality far more complex than the hypothesis suggests (McCalpin, 1996) . There seems to be segmentation on different scales. A case study in the Basin and Range province (normal and strike-slip faults) shows that small earthquakes (Mw < 7) rupture within their segment, whereas larger (Mw > 7) earthquakes produce more complex ruptures over multiple segments (dePolo et al., 1991).

Segmentation in subduction faults is apparent from the historical earthquake record. No known subduction earthquakes have ruptured along the entire subduction zone. In Chile, segmentation is well defined. Barrientos (2007) describes seven seismogenic zones along the contact beween the Nazca and South American Plate, and two more on the Antarctic Plate, (Fig. 4.5). Metzner et al. (2012) use corals from atolls along the Sunda megathrust off the coast of Sumatra, Indonesia, to study uplift during large earthquakes. These uplift events reveal that the boundary between the rupture zones of the megathrust earthquakes in 2004 and 2005 has been persistent for at least 1100 years. During this period at least 7 earthquake ruptures have terminated on this boundary and none have crossed it. Segment boundaries on thrust faults are often located at physical prominences like fracture zones, seamounts, oceanic plateaus or subducting plate boundaries (Carver & McCalpin, 1996).

Segments are, however, not necessarily ever persistent. Evidence from Nankai trough in Japan (1707) shows that some earthquakes can break through segment boundaries and rupture adjacent segments. The 1960 Valdivia earthquake, for example, was not confined to the Valdivia segment, but broke through the segment boundary and also ruptured part of the Concepción segment (Carver & McCalpin, 1996). The 2004 Sumatra earthquake and 2011 Tōhuku earthquake in Japan ruptured over an area that had only experienced small ruptures in modern history (Meltzner et al. 2012).

There might be a seismic cyclicality in many subduction settings (Carver & McCalpin, 1996). Lomnitz (2004) proposes a process of clustering, wherein periods of high seismic activity are preceded and followed by a period of low seismic activity. The debate is still hindered by insufficient and incomplete datasets.

In cases where multiple ruptures are known for a single segment in date and magnitude, a recurrence time can be calculated. This yields a time-predictable model for future earthquakes, by means of which the timing of the next earthquake of a certain size can be estimated (Nishenko, 1985). The Valparaíso segment, for example, exhibits a very constant recurrence time of 83 ± 9 years (Comte et al., 1986). By comparing the time since the last

22 large rupture with the average recurrence time, it is possible to estimate the seismic potential of a segment. A segment that has ruptured recently will have a low seismic potential. If the time since the last rupture exceeds that of the recurrence interval, the segment has a high seismic potential (Nishenko, 1985). Earthquakes tend to occur on segments with a high seismic potential (Sykes, 1971; in Lomnitz, 2004). This is the hypothesis of the seismic gap. For example, the Concepción segment was recognized as a seismic gap before the 2010 earthquake (Campos et al., 2002). Silbergleit & Prezzi (2012) have evaluated the seismogenic zones described in Barrientos (2007). The largest seismic gap in Chile at the moment is the Arica gap in the north of the country. There has not been a major subduction on this part of the fault since 1877, and it already has the potential of an earthquake with a moment magnitude of Mw 8.6.

3.3 Magnitude and intensity scales There is a need for a standardized, objective scale to express the greatness of an earthquake, and over the years several scales have been developed for this purpose. The scales can be divided into two families. The first group are the magnitude scales. They all have in common that they express the released energy of an earthquake. All magnitude scales are logarithmic scales, meaning that one step increase in magnitude equals a ten-fold increase in amplitude (McCalpin, 1996). The second group are the intensity scales. Instead of describing the released energy, this type of scale expresses the local effects of an earthquake.

3.3.1 Magnitude scales An overview of several magnitude scales can be found in McCalpin (1996). The first widely- used scale is the local magnitude (ML), better known as the Richter scale. It is still sometimes used in the media, but is has lost importance in scientific work. The Richter scale is accurate up to ML 6.5, but it saturates above this magnitude and progressively underestimates the released seismic energy (Fig. 3.1). Another problem is that this method was developed for measurements within 100 km of the epicenter.

The surface-wave magnitude Ms is based on the same principal as the Richter scale, but it uses amplitudes of low-frequency surface waves. The scales agree at Ms ≈ 6. It can measure earthquakes worldwide and does not require a seismograph within 100 km. For this reason many historic earthquakes are expressed in surface-wave magnitude (e.g. Lomnitz, 2004). It does, however, saturate above Ms 8 (Fig 3.1).

The currently most frequently used magnitude scale is the moment magnitude (MW). This scale is based on the seismic moment (MO) of an earthquake:

MO = D A µ

D is the average displacement over the entire fault, A is the area of the fault surface (the length of the fault zone multiplied with the depth of the fault zone), and µ is the average shear rigidity of the faulted rocks. Typical crustal rocks have a shear rigidity of 3.0 to 3.5 * 106

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N/cm2. The moment magnitude is calculated from the seismic moment using the following equation (Hanks & Kanamori, 1979):

Mw = 2/3 log MO – 10.7

The moment magnitude scale does not saturate at high magnitudes. Mw is equal to Ms for Ms

≤ 8 according to Kanamori (1983), or for magnitudes 6.2 ≤ Ms ≤ 8.2 according to Scordilis (2006) (Fig 3.1).

Figure 3.1 Relation between moment magnitude Mw (dashed line) and several other magnitude scales. ML and Ms. MJMA = Japanese Metrological Agency, MB = long period body wave, Mb = short period body wave. (Boore & Joyner, 1982) 3.3.2 Intensity scales A completely different concept are intensity scales. Those describe the amount of damage by an earthquake in an arbitrary manner on a certain location. The intensity depends mainly on the distance to the epicenter and the type of subsurface (Weishet, 1963; in Van Daele, 2013). Musson et al. (2010) give an overview of several intensity scales and their relationships. The original intensity scale was developed in the 19th century, and has been modified several times since. The most popular scales at present are the Modified Mercalli Intensity (MMI) scale, the Medvedev, Sponheuer and Kárník (MSK) scale, and the European Macroseismic Scale (EMS). The MMI and MSK scale are very similar to an extent that they can be used interchangeably. In this thesis intensities will be expressed on the MMI scale. The MMI scale is divided into 12 categories, labeled with a Roman numeral: (Wood & Neamann, 1931; in USGS, 2013b)

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I. Instrumental Not felt except by a very few under especially favorable conditions. II. Weak Felt only by a few persons at rest, especially on upper floors of buildings. Delicately suspended objects may swing. III. Slight Felt quite noticeably by persons indoors, especially on upper floors of buildings. Many people do not recognize it as an earthquake. Standing motor cars may rock slightly. Vibration similar to the passing of a truck. Duration estimated. IV. Moderate Felt indoors by many, outdoors by few during the day. At night, some awakened. Dishes, windows, doors disturbed; walls make cracking sound. Sensation like heavy truck striking building. Standing motor cars rocked noticeably. V. Rather strong Felt by nearly everyone; many awakened. some dishes, windows broken. Unstable objects overturned. Pendulum clocks may stop. VI. Strong Felt by all, many frightened. Some heavy furniture moved; a few instances of fallen plaster. Damage slight. VII. Very strong Damage negligible in buildings of good design and construction; slight to moderate in well-built ordinary structures; considerable damage in poorly built or badly designed structures; some chimneys broken. VIII. Destructive Damage slight in specially designed structures; considerable damage in ordinary substantial buildings with partial collapse. Damage great in poorly built structures. Fall of chimneys, factory stacks, columns, monuments, walls. Heavy furniture overturned. IX. Violent Damage considerable in specially designed structures; well-designed frame structures thrown out of plumb. Damage great in substantial buildings, with partial collapse. Buildings shifted off foundations. X. Intense Some well-built wooden structures destroyed; most masonry and frame structures destroyed with foundations. Rail bent. XI. Extreme Few, if any (masonry) structures remain standing. Bridges destroyed. Rails bent greatly. XII. Catastrophic Damage total. Lines of sight and level are distorted. Objects thrown into the air.

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4. Setting

4.1 Geographical Setting Chile is a long and narrow, north-south oriented country (4275 km length) along the Pacific coast of South America. It has a total surface of 756102 km². Because of its length parallel to the subduction zone, not even the largest earthquakes are felt in the entire country. The Great Chilean Earthquake of 1960, for example, had a rupture zone of more than 800 km (Kanamori & Cipar, 1974), which is less than a quarter of the total length of the country.

Chile has a population of approximately 17 million people, of which more than 6 million live in the capital Santiago and the surrounding area, known as the Santiago Metropolitan Region. This makes up over 35% of the total population (UNdata, 2013). The population density of Santiago averages 8470 people per km², whereas the country’s average is only 22,6 people per km².

The geographical landscape can be divided into three parallel north-south oriented zones. Along the coastline of Chile lies the Chilean Coastal Range or Cordillera de la Costa. This mountain range is separated from the Andes by a depression called the Chilean Central Valley or Valle Central. The capital city of Santiago is located in this valley. East to the central valley lie the Andes mountains, the world’s second highest mountain range.

Because of the north-south orientation of the country and its mountainous terrain, Chile contains several climate zones. The northern zone, or Norte Grande, between 17° S and 27° S is a desert region with very low annual precipitation. The Atacama desert lies in this region and is the driest desert in the World. Between 27° S and 32° S lies the Norte Chico region, which has a semi-arid climate, with precipitation during the austral winter. The climate of the Central Zone (32° S to 38° S) is characterized by a Mediterranean climate with cold, wet austral winters and dry austral summers. The Santiago Metropolitan Region, the research area in this thesis, lies in the Central Zone. The mean annual precipitation in Santiago is approx. 550 mm (von Gunten et al., 2009a), but there are strong climate fluctuations. Those fluctuations are mostly related to the El Niño-Southern Oscillation (ENSO), which causes alternations between years with high precipitation (El Niño years), years with low precipitation (La Niña years) and normal years (Meza, 2005). The prevailing wind direction is south to south-west (Grass & Cane, 2008). The South Zone lies between 38° S and 42° S. It is known for its humid and temperate climate. The last climate zone is the Austral Zone between 42° S and 56° S. It has a cold climate with constant precipitation throughout the year (Muñoz et al., 2007).

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Figure 4.1: Map of South America. The first inset shows the Santiago Metropolitan Region. The second inset shows the location of the lake. LLE = Lago Lo Encañado, LN = Laguna Negra, EEY = Embalse El Yeso. Satelite pictures an aerial photographs from Google Earth. Lago Lo Encañado, also known as Laguna del Encañado, is a small proglacial lake situated in the Andean mountains, 50 km ESE of Santiago (33°40’S/ 70°08’W; Fig. 4.1). The location of the lake, in the proximity of the capital Santiago, adds to the importance of studies like this. Santiago is the most densely populated area of Chile, and any earthquake in this region will increase the risk of casualties and damage in comparison to earthquakes in the lesser populated areas. It is therefore important to know the seismic history of this region.

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Figure 4.2 Photo of Lago Lo Encañado, taken from the eastern shore. On the right is the Rio del Encañado, forming a delta. On the left is the outlet towards the Maipo river. Photo by P. Kempf. The lake has a length (N-S) of approximately 900 m and a width (E-W) of 650 m, with a surface area of 0.5 km2. Its maximum depth is approx. 35 m. Lago Lo Encañado has a very large catchment area compared to the size of the lake. The catchment (approx. 30 km2) is 77 times larger than the lake (Salvetti, 2006) (Fig. 4.3). The catchment area is a glacial, mountainous region, producing a high amount of eroded, clastic material. To the north, the Río Lo Encañada flows through the glacial valley into the lake, forming a delta. The glacier, Echaurren Norte, is situated in the most distal part of the catchment. This glacier shows strong interannual mass- balance variations due to its sensitivity towards ENSO events, but in general the mass-balance has decreased since 1988 (Casassa et al., 2006). The delta contains a large amount of plants and algae. The western and eastern flanks are both very steep, starting directly from the lake shore. About 200 vertical meters above the eastern flank of the lake lies Laguna Negra. The water from this larger lake used to run into Lago Lo Encañado, which produced two colluvial fans to the northeast and southeast. There are several smaller fans on the western flank. Originally, water would leave the lake to the south and flow via the Estero del Manzanito and the Maipo river into the Central Valley, towards the suburbs of Santiago. The flow of water from Lago Lo Encañado, Laguna Negra and the artificial reservoir Embalse El Yeso is now controlled by man-made dams. The Santiago Metropolitan Region lies in the Central Zone, which is known for its dry summers. Water reserves are built in the Andes to regulate the water supply for the city. At a height of 2490 m above sea level, the lake lies above the tree- line. The landscape is barren with a low sub-alpine shrub-land environment.

Lake deposits are normally finely laminated, consisting mainly of fine-grained sediments. The average sedimentation rate is 1.75 mm/year (Salvetti, 2006), but the amount and type of

29 sediment changes seasonally and annually. This is due to several factors, including changes in precipitation and glacier melting. The laminations are sometimes disrupted by deposits that are often coarser in grainsize and less structured.

Figure 4.3 Satellite image of Lago Lo Encañado (Google Earth). The green line delimits the catchment of the lake. The red line indicates a change in the geology of the catchment. The grey striped area is the current location of the Echaurren Norte glacier. The blue line is the Rio del Encañado and its main tributaries.

4.2 Geological setting

4.2.1 Tectonic Setting The western coast of South America has been an active plate boundary for most of the geological time. The oldest known tectonic cycle happened between the Late Proterozoic and the Early Cambrian. Only during the latest Permian to the earliest Jurassic the subduction was interrupted. The Andean stage of subduction started in the Early Jurassic and is still active

30 today (Charrier et al., 2007). It is part of the Ring of Fire, a term used to describe almost the entire coastline of the Pacific Ocean, running from New Zealand via the eastern coasts of Asia and the North-American west coast all the way to the south of Chile. All along this ring oceanic crust is being subducted under continental crust. This produces large volcanoes and frequent eruptions, hence the name. In Central Chile it is the oceanic Nazca Plate that subducts underneath the South-American Plate at a rate of 66 mm/year (Cembrano & Lara, 2009) to form the Andes mountain range. In the south of Chile the oceanic Antarctic Plate is subducted below the South-American Plate. The Antarctic plate moves at a lower rate, about 1-2 mm/ year (Dietrich et al., 2004). The plate boundary between the Nazca Plate and the Antarctic Plate is formed by a spreading ridge called the Chile ridge. The Chile ridge collides with the South-American Plate at a latitude of approx. 46° S. The point where the three plates are in direct contact is called the Chile Triple Junction or CTJ (Maksymowicz et al., 2012) (Fig 4.4).

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Figure 4.4 Terrain map of south Chile. The spreading ridge can be seen in the bathymetry. The Nazca Plate (north) and the Antarctic Plate (south) are subducted below the South American Plate. CTJ = Chile Triple Junction. The red arrows show the movement of the two plates. Adapted from Maksymowicz et al., 2012. When an oceanic plate subducts under a continental plate, the entire oceanic plate has to move at the same average rate. However, the oceanic and continental plates are locked to a certain extent in the part involved in the subduction. The relative movement between the two plates in this part is accommodated by the interpolate thrust fault in the form of earthquakes (Carver & McCalpin, 1996). Two end members are described in Plafker (1972; In Carver & McCalpin, 1996). In Marianas type subduction zones, the subducting plate consists of older, and therefore heavier, oceanic crust. Typical for this type of subduction is a high dip angle and frequent but moderate seismic activity. On the other hand, Chilean type subduction zones are areas where young, lighter, oceanic crust is subducted. The coupling between the two plates is stronger, the angle under which the plate is subducted is lower and the earthquakes are larger but less frequent. Evidently, the Chilean coast is a Chilean type subduction zone.

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Historic earthquake records from Central and South-Central Chile reveal three distinct segments. From North to South, these are: i) the Valparaíso segment (ruptured in 1985), ii) the Concepción segment (where the 2010 earthquake occurred) and iii) the Valdivia segment (where the 1960 earthquake ruptured) (Fig. 4.5). The Valparaíso segment is subducted below the Santiago Metropolitan Region. Lago Lo Encañado will therefore be influenced mostly by earthquakes occurring on the Valparaíso segment, and to a lesser extent by earthquakes on the Concepción segment.

Figure 4.5 a) An overview of the major earthquake segments along the western South-American coast. The red circle indicates the location of Santiago. b) A closer view on the Valparaíso, the Concepción and the Valdivia segment, together with the rupture zones of the most important historical earthquakes on these segments. The epicenter of the 1960 Valdivia earthquake is indicated. Adapted from Melnick et al., 2009. Subduction earthquakes, or megathrust earthquakes, are among the strongest types of earthquakes, because they generally rupture large fault segments. In addition to these megathrust earthquakes, the process of subduction causes several other geohazards. Volcanoes are formed all along the subduction zone. Some authors believe that earthquakes can trigger volcanic eruptions. Lara et al. (2004) link a fissure eruption in the Cordón Caulle volcanic complex to the 1960 Valdivia earthquake. Landslides, or lahars, triggered by loading of volcanic ashes or by earthquakes also cause many casualties, for example on the slopes of the Nevado del Ruiz in Colombia (Stern, 2004). The stress on the South-American Plate also produces intraplate earthquakes. These earthquakes are usually weaker and shallower in

33 comparison to the large megathrust earthquakes, but can locally be as devastating. The largest ever recorded intraplate earthquake in Chile was the Chillán earthquake (Mw 7.8) on 25 January 1939 (Madariaga et al., 2010), the deadliest earthquake in the history of this country with a death toll of more than 28.000 (USGS, 2012a).

4.2.2 Local Geology Lago Lo Encañado is situated in a region where the geology consists of volcanic and volcanoclastic deposits of Oligocene to Miocene age and Cenozoic plutonic basement (SERNAGEOMIN, 2003) (Fig. 4.6). The volcanic rocks are of basaltic to andesitic or basaltic nature. There are four units in the proximity of the lake. Lago Lo Encañado lies in the unit OM2C. North of the lake, within the catchment area, unit M3i outcrops. To the south of Lago Lo Encañado lies an outcrop of unit PPl1r, and north of Laguna Negra unit Msg reaches the surface. These last two units are not found in the catchment of the lake. The explanation of these codes was translated from SERNAGEOMIN (2003).

OM2C: This unit consists of volcanosedimentary sequences. It contains lava that is basaltic to dacitic in composition, together with epiclastic and pyroclastic rocks. The age of the rocks is Oligocene to Miocene.

M3i: This unit is made up of Lower to Middle Miocene volcanic complexes that are partly eroded. The volcanic sequences consist of lavas, breccias, domes and pyroclastic deposits. The composition is andesitic-basaltic to dacitic.

PPl1r: The unit south of Lago Lo Encañado contains landslide deposits of Pliocene to Pleistocene age. The landslide deposits consist of polymictic breccias with a matrix of sand or silt in varying proportions.

Msg: This unit is Late Miocene in age (13-7 Ma). It is mainly made up of hornblende- biotite granodiorite, and to a lesser extend of monzogranites, quartz bearing monzonites and monzodiorites.

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Figure 4.6 Geological map of central Chile. the westernmost lake in the red box is Lago Lo Encañado. The lake is enlarged for visibility purposes. Adapted from SERNAGEOMIN (2003). See text for description of lithological units OM2C, M3i, PPl1r and Msg.

The area around Lago Lo Encañado experiences some volcanic activity as it is situated in the northern part of the Southern Volcanic Zone (SVZ). The SVZ is the most active of the three volcanic zones in Chile, with an average of one eruption per year (Stern et al., 2007). The other volcanic zones are the Austral Volcanic Zone further to the south and the Central Volcanic Zone in northern Chile. There are at least three active volcanoes in the direct proximity of the lake (Fig. 4.7). The closest volcano is the San José volcano. It last erupted in 1960, and shows permanent solfataric activity. To the north is the Tupungatito Volcano which last erupted in 1987. It also shows permanent solfataric activity. The Maipo Volcano is located south of the San José Volcano. It last erupted in 1912. It lies within a larger caldera called the Diamante Caldera. These volcanoes lie on a parallel north-south line east of Lago Lo Encañado.

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Figure 4.7 Aster satellite elevation map of the Santiago region. LLE = Lago Lo Encañado. The location of the three closest volcanoes is indicated. The lake is enlarged for visibility purposes. The blue lines represent the Maipo river and its most important tributaries.

4.3 Historical setting The first evidence of human settlement in what is now the Republic of Chile was found at the archeological site of Monte Verde, near the town of Puerto Montt, and dates back to at least 12.500 years ago (Dillehay, 1999). However, the maximum age is still under debate (Dillehay, 1999; Dickinson, 2011). Since then, the country has been inhabited by different indigenous peoples, of which the Mapuche were the most numerous. The Mapuche had no written language. This is very unfortunate, because it means that there is no historical record of earthquakes in pre-Hispanic ages. This limits accuracy of statistical predictions for future earthquakes. The European discovery of Chile was led by Ferdinand Magellan on 1 November 1520, but the conquest of Chile happened in 1541, when Pedro de Valdivia led a Spanish army from Peru in an attempt to extend the Spanish Empire southward. He founded the town of Santiago de Nuevo Extramadura, now Santiago de Chile. The southern border of the Empire was the Bío Bío River near the city of Concepción, where the Mapuche set up their defense line. This means that there is no written evidence of earthquakes before 1541 from Northern and Central Chile, and sparse evidence from Southern Chile before the 18th century. But even after the arrival of the Spanish, Chile remained scarcely populated. For example, the population of Santiago was around 5000 during the 1647 earthquake (Udías et al., 2012). Still fewer people lived in the rural areas around Santiago or along the coast. The coastal town of Valparaíso is described as “a humble hamlet at best” in 1580 (Cisternas et al., 2012). The amount of written evidence of earthquakes is therefore limited. Most research on this subject is based upon two important studies of historical earthquakes in Chile, composed by Perrey (1854; In Cisternas et al., 2012) and Montessus de Ballore (1911-1916; In Cisternas, 2008). The most complete overview of earthquakes and tsunamis in Chile between 1535 and 1960 was written by Lomnitz (1970b; 2004). This overview is, however, not perfect. Cisternas et al.

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(2012) found evidence of a large earthquake in 1580, which is not mentioned in previous studies. Other authors who studied the primary historic sources include Udías et al. (2012), who have studied the Archivo General de Indias in Seville. This archive preserves documents from Spanish colonial administration, including letters that describe the damage of several large earthquakes.

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5. The historical archive of major earthquakes in Central Chile

This chapter summarizes the historical earthquakes in Central Chile with potentially high seismic intensities in the Santiago Metropolitan area. It includes both large megathrust earthquakes and local intraplate earthquakes, which are discussed separately. The subduction earthquakes are further divided per segment, from North to South: the Valparaíso segment, the Concepción segment and the Valdivia segment (Fig. 4.5). All seismic intensities are expressed on the Modified Mercalli Intensity scale unless otherwise mentioned.

5.1 Subduction earthquakes

5.1.1 Valparaíso Segment rd On the 3 of March 1985, an interplate earthquake with Ms 7.8 (Barrientos, 1988) struck the coast off Valparaíso (Fig. 5.1). The rupture zone had a length of more than 140 km between approx. 33⁰S and 34⁰S (Fig. 4.5). In the Argentinean city of Mendoza a seismic intensity of VI was reported. This earthquake was also forecasted based on constant repeat times (83 ± 9 years) of earthquakes on this segment (i.e. Valparaíso segment; Comte et al., 1986). Shortly before the earthquake, Nishenko (1985) saw the Valparaíso segment as the segment with the highest seismic potential, and predicted an earthquake within twenty years (1984-2004).

The earthquake on the 6th of April 1943 had an epicenter close to the town of Illapel (Fig. 5.1).

It had a Ms of 8.3 (Lomnitz, 1970b; Beck et al., 1998). The rupture zone is estimated between 30.2⁰ S and 32.2⁰ S (Kelleher, 1972). The earthquake produced a local tsunami with wave heights of up to 4 m and killed eleven people. The earthquake was felt as far as the Argentinean capital of Buenos Aires. Despite the high magnitude this earthquake has not received too much scientific attention and seems to be forgotten today (Lomnitz, 1970a).

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Figure 5.1 Map of Central Chile with the locations of cities and rivers mentioned in connection with earthquakes on the Valparaíso segment. LLE = Lago Lo Encañado, enlarged for visibility purposes.

On 17 August 1906 Valparaíso was hit by a large interplate earthquake. This is a peculiar earthquake as it happened within 30 minutes after a major earthquake in the Aleutian Islands (Okal, 2005). The few seismic stations that existed at that time had major difficulties in separating the seismic waves from both earthquakes. Therefore the position of the epicenter, the rupture zone and the magnitude of the Chilean earthquake are estimated with a high uncertainty. However, it is estimated that the location of the rupture zone was similar to the

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1985 Valparaíso earthquake, with a rupture zone of 365 km between 31.7⁰ S and 35⁰ S (Comte et al., 1986), and an Ms of 8.3 (Kelleher, 1972). It was long thought that the tsunami that hit Hawaii with run-up heights of up to 3.5 m originated from the Aleutian earthquake (Lomnitz, 1970b). However, a re-evaluation of both earthquakes showed that the Aleutian event was an intraplate earthquake, and that it was the Chilean earthquake that produced the tsunami (Okal, 2005).

An earthquake struck Valparaíso on the 19th of November 1822. It produced a moderate tsunami with maximum wave heights of about 4 m. The seismic intensity felt in Santiago was between VII and VIII, inferred from damage to public buildings. The town of San José de Maipo in the Maipo Valley was not damaged by this earthquake (Lomnitz 1970b). The magnitude is estimated between Ms 8 and 8.5 (Lomnitz, 2004). There was another earthquake

(Ms 7) in the proximity of the 1822 epicenter in 1829, which was possibly a late aftershock (Lomnitz, 2004).

On the 8th of July 1730 central Chile was hit by a severe earthquake. An area with a length of more than 1000 km (from Copiapó to Concepción) was affected. Valparaíso and Concepción were destroyed by a tsunami, which reached the coast of Japan. The damage in Santiago led to a seismic intensity estimated between X and XI (Udías et al., 2012). The death toll was limited because people were warned by a moderate foreshock. Montessus de Ballore (1912; in Udías et al., 2012) describes this earthquake as the largest earthquake that ever occurred in Chile before 1912 (the date of publication). The extent of the damage supports the statement of Montessus de Ballore and indicates a magnitude of Ms 8.5 to 9.0 (Lomnitz, 2004).

On 13 May 1647 the city of Santiago was hit by a large earthquake that killed one-fifth of its population (1000 people). The shock was felt between the Aconcagua River and the Maule river over a north-south length of 420 km. The epicenter must have been within 50 miles of Santiago at most (Lomnitz, 1970b). The assumed epicenter near Quillota (after Montessus de

Ballore; in Lomnitz, 1970b) lies within this radius (Fig. 5.1). The Ms of the earthquake is estimated to be between 8 and 8.5, based on the extent and the duration of the shaking. An origin of the earthquake cannot be inferred with certainty, but a subduction setting, similar to the 1906 Valparaíso earthquake is likely (Lomnitz, 1970b; Udías et al., 2012).

5.1.2 Concepción Segment The latest major earthquake on this segment is the 2010 Maule earthquake (Fig. 5.2). On the 27th of February 2010 this earthquake ruptured a zone between 34⁰S and 38⁰S on the

Concepción segment (Wang et al., 2012). The earthquake had a Mw of 8.8. The surface-wave magnitude was Ms 8.5. At least 523 people were killed by the earthquake and the Pacific-wide tsunami generated by it (USGS, 2013c). This region was recognized as a mature seismic gap since the 1939 earthquake had been reinterpreted as a large (Mw 7.8) intraplate earthquake (Madiaraga et al., 2010). There had been no major megathrust earthquakes on this segment since the 1835 earthquake (Campos et al., 2002). The rupture area had a length of approx. 400

41 km and a width of 140 km (Pulido et al., 2011), and an average slip of 8.1 m releasing most of the accumulated stress on the segment since the 1835 earthquake (Wang et al., 2012).

Figure 5.2 Map of Central Chile with the location of cities on the Concepción segment.

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On 20 February 1835 Concepción was hit by an earthquake with an estimated Ms between 8 and 8.5 (Lomnitz, 2004). Darwin (1845) describes the city of Concepción as being completely in ruin after the earthquake and reported traces of a tsunami with a “height of 23 vertical feet [approx.. 7 m] above the highest springtides”. Lomnitz (1970b) mentions no damage in Santiago, but he warns that “Chilean earthquakes tend to be underrated in their effects, because of a national posture of understatement that is most apparent in journalism”. The same author corrects this statement in a later publication, mentioning “damage in Santiago” (Lomnitz, 2004).

On the 25th of May 1751 a large earthquake hit the city of Concepción. A large tsunami swept through the city and caused damage throughout the Pacific Ocean. The population was warned by at least two foreshocks just before the main shock. To avoid similar destructive tsunamis in the future, the population of Concepción decided to relocate the city to higher ground. Unfortunately the new city was located on an area of soft sediment, which increased the damage during the 2010 earthquake (Udías et al., 2012). There is no mention of the precise rupture zone, but it is described as being similar to that of the 2010 earthquake. The damage to buildings in Santiago suggests an intensity of the order VI to VII on the MMI scale.

The magnitude is estimated as Ms 8.5 in Lomnitz (1970b and 2004), a value that Udías et al. (2012) describe as “conservative”.

The earthquake on the 15th of March 1657 and the following tsunami completely destroyed the town of Concepción. The epicenter is thought to have been offshore, which is in agreement with the tsunami generation (Lomnitz, 1970b; Lomnitz, 2004; Udías et al., 2012). Concepción was the southernmost stronghold of the Spanish Empire, and there is no information on this earthquake further south. The shaking definitely extended up to the Maule river to the north (170 km). Udías et al. (2012) quote a letter from 1657 saying that everything in Santiago that was rebuild after the devastating 1647 earthquake (see below) was again destroyed. Lomnitz (1970b) doubts these historic sources, claiming that the earthquake was barely or not at all felt in Santiago. They agree, however, that this earthquake was smaller than the 2010 earthquake and similar to the 1835 earthquake that all ruptured within the same segment. The magnitude is estimated as Ms 8 or slightly less (Lomnitz, 2004).

The oldest known historical earthquake in central Chile occurred on the 8th of February 1570. A large earthquake, followed by a tsunami, hit the town of Concepción. Based on reports of damage to houses and of cracks in the ground, this earthquake is thought to have had a larger st magnitude than the earthquake on the 21 of May 1960 (Ms between 7.5 and 7.9), but nd smaller than the earthquake on the 22 on May 1960 (Ms 8.3). Therefore a magnitude of approx. Ms 8 is estimated (Lomnitz, 1970b).

5.1.3 Valdivia Segment On the 22nd of May 1960, the coast of South-Central Chile was hit by an earthquake with an

Mw between 9.5 (Kanamori, 1977; in Lara et al., 2004) and 9.6 (Lomnitz, 2004), and an Ms of 8.3 (Kanamori & Cipar, 1974).. This is to date the strongest earthquake ever recorded

43 worldwide. This earthquake ruptured the entire Valdivia segment, and had a rupture length of more than 800 km between Concepción in the north to the Chile Triple Junction in the south (Plafker & Savage, 1970; Barrientos & Ward, 1990)(Fig. 5.3). The earthquake and the generated tsunami cost approx. 1655 lives in Chile, and another >200 along the coasts of the Pacific Ocean (USGS, 2012b). It was preceded by several large foreshocks on the 21st and 22nd of May, with magnitudes of up to Ms 7.9. These foreshocks forced people to leave their houses, which explains the relatively low death toll. The same segment ruptured before in 1837, 1737 and 1575 (Lomnitz, 2004). Many of these earthquakes on the Valdivia segment in South Chile have caused severe damage in Central Chile, but mostly through tsunamis in the coastal villages. As the northern boundary of the segment is more than 500 km south of Lago Lo Encañado, even the largest earthquakes along the Valdivia segment are not expected to leave an imprint in this lake. The seismic intensity reached a maximum of V in Santiago. This was caused by the foreshock on the 21st of May which ruptured closer to Santiago than the main shock. The main shock produced seismic intensities of up to IV in the capital (Lazo, 2008). Therefore this study will not go any deeper into earthquakes along the Valdivia segment.

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Figure 5.3 Map with the location of the most important geographical features on the Valdivia segment. CTJ = Chile Triple Junction.

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5.2 Intraplate earthquakes th The 1958 Maipo Valley earthquake (Mw 6.3) struck on the 4 of September. Some authors refer to this earthquake as the Las Melosas earthquake (Alvarado et al., 2009). This earthquake cannot be described as a major earthquake in the seismic history of Chile, but it is an important earthquake in the history of the Maipo Valley, in which Lago Lo Encañado is situated (Fig. 5.4). According to Flores et al. (1960) three separate shocks were recorded th within 5 minutes, with Ms between 6.7 and 6.9. The earthquakes of the 4 of September formed the peak of a month of increased seismic activity. The maximum intensity in the valley (more precisely in the town of El Volcan) was estimated as high as IX-X on the MMI scale (Alvarado et al., 2009) (Fig. 5.4; Fig. 9.6). It did not cause damage in Santiago (Lomnitz, 1960). There have been local earthquakes before in 1850, 1870-80, 1883, 1905 and 1947 (Flores, 1960), but as there is no additional information to be found on these earthquakes (except for the 1850 earthquake), they will be regarded as minor events. Lomnitz (1960) reports another earthquake in 1945 in the valley, which did not cause any damage.

The Chillán earthquake on the 25th of January 1939 is the deadliest earthquake with a death toll of about 30.000. The epicenter was northwest of Chillán (Fig. 5.2), and the estimated Ms differs depending on the authors between 7.8 (Leyton et al., 2009) and 8.3 (Lomnitz, 1970b). This earthquake was long thought to be an interplate thrust earthquake (Nishenko, 1985), but is now recognized as an intraplate earthquake on a normal fault (Beck et al., 1998). This matches with the lack of tsunamis, the inland location of the epicenter and the high seismic intensities.

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Figure 5.4 A map showing the most important locations in the Maipo Valley. The blue line represents the Maipo river and its most important tributaries. LLE = Lago Lo Encañado, LN = Laguna Negra, EEY = Embalse El Yeso.

In 1861 an earthquake hit Mendoza, Argentina, killing more than 6000 people (Lomnitz, 1970a), or up to 12000 in other studies (Rickard, 1863; in Perucca & Moreiras, 2006) (Fig. 5.1).

The Ms is estimated at 7.2 (Perucca & Moreiras, 2006). The epicentre of the earthquake was situated in the western suburbs of the city of Mendoza (Silgado, 1985).

On the 6th of December 1850 an intraplate earthquake hit the Maipo Valley. The seismic intensity reached in Santiago was VII on the MMI scale. The epicentre is assumed to be similar to the 1958 Maipo Valley earthquake, based on the reports of large rockslides 14 km south of San José de Maipo (Fig. 5.4). The magnitude is estimated between Ms 7 and 7.5 (Lomnitz, 1970b).

5.3 Earthquakes with an unknown source A recent discovery in the Spanish archives contains evidence for a large earthquake in the Santiago area on the 7th of August 1580 on the Julian calendar (Cisternas et al., 2012). The most important conclusion of this discovery is that the historical archive of Central Chile is incomplete. Most studies are based on the works of Perrey and Montessus de Ballore, and few recent studies look at the primary sources. The earthquake must have had a higher

47 seismic intensity than the 1575 earthquake in Santiago, because the damage was significantly larger. The intensity is estimated as greater than VII, based on comparison of damage to similar buildings after recent earthquakes. No casualties were reported. There is no knowledge of the source of the earthquake. The possibility of an intraplate earthquake with an epicenter in the Andes, or a deep hypocentre on the downgoing oceanic Nazca crust is excluded after comparison the effects of the 1580 earthquake to that of the 1958 Maipo Valley and the 1939 Chillán earthquake respectively (Cisternas et al., 2012). A subduction earthquake on the Valparaíso segment, similar to the 1730, 1822, 1906 and 1985 earthquakes is a viable possibility, and fits in the time frame of recurrence on this segment. The large number of small aftershocks is similar to the situation following the 1985 earthquake (91 recorded aftershocks; Comte et al., 1986). Another possible rupture mode of this earthquake is a shallow intraplate earthquake with an epicenter in the proximity of Santiago. There are no reports of a tsunami after this earthquake. This could be an argument in favour of a local intraplate earthquake, if it is not explained by the very small population living along the coast in the 16th century. No magnitude is mentioned.

The same region was hit by another earthquake on the 17th of March 1575. The intensity in Santiago was between VI and VII on the MMI scale. No deaths were reported (Lomnitz, 2004). The damage was smaller than the damage of the 1580 earthquake (Cisternas et al., 2012). Similarly, the understanding of this earthquake is limited because of the fact that the region around Santiago was scarcely populated. Lomnitz (1970b) sees the epicenter in the proximity of La Ligua, a town at the coast east of Santiago with an estimated Ms of 7 to 7.5. An epicenter close to the coast could indicate an interplate source. Comte et al. (1986) see this event as the predecessor of the earthquakes in 1730, 1822, 1906 and 1985 on the Valparaíso segment. This would be consistent with the average return of large earthquakes (83 ± 9) on this segment. No aftershocks were mentioned in the Spanish archive, nor was there mention of a tsunami (Lomnitz, 2004).

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6. Previous studies on Lago Lo Encañado and nearby lakes

6.1. Lago Lo Encañado Lago Lo Encañado was chosen as one of the lakes in which evidence of the 2010 Maule and 1960 Valdivia earthquake was sought by Van Daele (2013). Cores ENC01 and ENC02 were used in that work. Van Daele (2013) demonstrates the presence of the 2010 event and the absence of the 1960 event in these cores.

One other scientific study has been performed on Lago Lo Encañado. It is a paleolimnological case study on the lakes of the South-Central Andes in the context of an MSc thesis by Salvetti (2006), at the University of Bern, Switzerland. It included one gravity core, labeled core 1305. The core is taken in the center of the lake (Fig. 7.1). The core was not processed during the present study, but it is important because this core contains dated layers (Fig. 6.1). It will therefore be described shortly. Six coarse sediment layers were recognized. Unstructured (homogenites) or deformed sedimentary units can be found below the larger coarse sediment layers. There is a small layer with an abundance of organic matter. The rest of the core consists of “laminated to finely banded sediment”. A 14C dating has been performed on two layers. The layer at 21.5 cm yielded an age of 1002 ± 19 AD, and that at 29.5 cm was dated at 593 ± 32 AD. It is however unclear which kind of material was used for dating. In the methods section there is mention of organic macrofossils. In the results chapter it says that these layers were dated with organic microfossils. The upper 14 cm have been dated by 210Pb. The homogenite (16-21 cm; Fig. 6.1.) between the end of the Pb dating and the first 14C date represents a hiatus of approx. 900 years. The average sedimentation rate above the hiatus is 7 years per centimeter, or 14 mm per year. Under the hiatus the average sedimentary rate is 50 years per centimeter, or 0,2 mm per year. This is at least partly caused by compaction, but cannot explain a difference of an order in magnitude. The 210Pb dating proved to be very useful in this study. There were some problems involving the 14C data. The first problem is that one of the dates (the upper date at 21.5 cm) was taken in a turbidite. This means that the age of 1002 AD is a maximum age, and all we know is that the turbidite is younger than the proposed age. The other date was taken in a part of background sedimentation, but also has an extremely high age. As there is no 14C dating in a section that has been dated by 210Pb and no clear mention of a working method we do not know if there is a reservoir age that has or has not been accounted for.

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Figure 6.1 Core 1305 in normal colours (left) and after histogram equalisation (2nd from the left). The core is logged (2nd from the right). The result of the two dating methods (210 Pb (upper 14 cm) and 14C) is combined into an age-depth model (right), adapted from Salvetti (2006).

6.2 Surrounding lakes Another local study that is used in this thesis is the work of von Gunten et al. (2009a, 2009b, 2009c). These authors have taken cores in five lakes in the Santiago Metropolitan Region to evaluate the amounts and effects of anthropogenic pollution. In four of these lakes, the position of turbidites is indicated in the age-depth model for the twentieth century. These lakes are Laguna Negra (immediately to the east of Lago Lo Encañado), Laguna del Inca (approx. 70 km north of Lago Lo Encañado in the Andes), Laguna El Ocho (south-west of Lago Lo Encañado in the Andes) and Laguna Aculeo (south of Santiago in the Central Valley) (Fig. 6.2).

Vandenberghe (2012) took several cores in Laguna Negra in the context of a paleoclimatological study. (Fig. 6.2). The position of a total of six possible seismites was described, dated and linked to historic earthquakes.

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Figure 6.2 Map of the Santiago area, showing the locations of Laguna del Inca (b), Laguna Negra (c), Laguna el Ocho (d), Laguna Ensueño (e) and Laguna Aculeo (f). The wind rose in the center of the map shows the predominant wind direction (Grass & Cane, 2008). (von Gunten et al., 2009a)

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7. Methods

The sediment cores studied in this thesis were collected during two field campaigns (January 2011 and February 2012). In the first campaign, two gravity cores were retrieved (ENC01 and ENC02; Fig. 4.1, Table 1), by M. Van Daele, W. Vandoorne and A. Peña. In the second field season, 7 more cores (ENC03-ENC09; Fig. 7.1, Table 7.1) were retrieved, by M. Van Daele, P. Kempf and W. Vandoorne. During this second field season also a sidescan sonar survey was performed (Fig. 8.2).

7.1 Coring All cores are gravity cores, taken by controlled free fall from a zodiac with a gravity corer. Plastic liners with a diameter of 6,3 cm were dropped with additional weight to penetrate the sediment. The cores were then winched up, sealed with caps, labeled and subsequently shipped to Belgium, where they were finally opened. Core ENC04 has not been opened for this thesis. It was taken at the same location as ENC03, in an attempt to retrieve a longer core. However, the penetration was not as deep as hoped, and therefore the core was not opened.

Table 7.1: Core lengths and locations, and year of coring. Core ENC04 was not opened.

Core ID Length (cm) Latitude Longitude Year ENC01 63 33°40'14.20"S 70° 8'4.09"W 2011 ENC02 87 33°40'19.60"S 70° 7'59.48"W 2011 ENC03 84,3 33°40'20.82"S 70° 7'58.91"W 2012 ENC04 #### 33°40'20.68"S 70° 7'58.91"W 2012 ENC05 43,2 33°40'20.35"S 70° 8'2.98"W 2012 ENC06 68,7 33°40'14.99"S 70° 8'6.07"W 2012 ENC07 55,3 33°40'13.08"S 70° 8'1.00"W 2012 ENC08 73,4 33°40'11.14"S 70° 7'55.99"W 2012 ENC09 102,3 33°40'20.71"S 70° 7'52.97"W 2012

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Figure 7.1 Satellite image of Lago Lo Encañado showing the coring locations. Note that ENC03 and ENC04 are taken at the same location. 1305 is the approximate location of the core analyzed by Salvetti (2006). 7.2 Geoacoustics 7.2.1 Sidescan sonar A sidescan-sonar survey is a relatively fast and easy way to map the sea- or lake floor. A boat tows the sonar fish through the water. The source emits an acoustic signal in two fan-shaped lobes perpendicular to the heading of the boat. The frequencies of this signal range between 1 kHz (wavelength = 1.5 m) and 1 MHz (wavelength = 1.5 mm). The range of the pulse decreases with a higher frequency, but it yields a better resolution, as the wavelength dictates the scale of objects that can be detected (Blondel, 2009). The smaller the wavelength (and simultaneously the higher the frequency), the higher the attenuation.

Lago Lo Encañado has a maximum depth of approx. 35 m, allowing a relatively high frequency while towing the sonar near the lake surface. The fish was towed approx. 0.5 m below the lake surface. This method has the advantage that bathymetrical data could be retrieved (Fig. 8.2). The sidescan sonar was towed by a zodiac in predominantly NNW-SSE directed survey lines (i.e. parallel with the lake axis) with overlapping swaths. The data were acquired using a Klein 3000 Side Scan Sonar, at a frequency of 500 kHz. The frequency of 500 kHz used in this thesis means that the emitted acoustic signal has a wavelength of 3 mm. The signal propagates

54 through the water column and is scattered when it hits the basin floor or an object. A part of the signal is reflected back toward the imaging sensor. The amplitude of the backscatter contains information about the nature of basin floor. This includes the direction of the slopes and the roughness of the surface. The bottom track provides information on the water depth below the fish. The width of the bottom track is inversely proportional to the depth. Unlike multibeam data, sidescan-sonar data only contain depth information of the line below the fish. The sidescan-sonar survey was performed under unfavourable circumstances during the 2012 field campaign, by M. Van Daele, P. Kempf and W. Vandoorne. The quality of the images is therefore not as good as it could be. It was not possible to perform a multibeam or seismic survey because of the remoteness and the protected status of the lake.

7.2.2 Sonarscope Sonarscope is a software developed by IFREMER in Brest, France, for the analysis of multibeam and sidescan-sonar data. The following steps were used for processing the data. First, the heading of the lines was plotted (Fig. 7.2). To improve the quality of the signals they were preprocessed, correcting for heading and artifacts. Small-frequency variations in the heading were filtered out, and artifacts were deleted by computing the bottom track, which correlates to water depth. Subsequently the files were merged to create a reflectivity map. The high reflectivity values of the slopes overwhelmed the more subtle reflectivity changes in the center of the lake. Because of this great difference in reflectivity, the data was processed separately. The lines where the boat turned were discarded, because the signal was too disturbed. Unfortunately, due to this cleaning operation, important information on the outer part of the lake was lost. The images of the center part of the lake, the northern edge and the western edge were exported to Google Earth and an interpretation map was created in Global Mapper (Fig. 8.5).

The bathymetry was mapped by extracting the depth layer in the sidescan-sonar data. This depth layer gives the depth along the towline. It was plotted together with the heading in Surfer. The bathymetry was interpolated between the survey lines.

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Figure 7.2 Sidescan-sonar survey lines with heading.

7.3 Core opening The cores were opened in the core opening laboratory at the Sterre campus of Ghent University (Department of Geology and Soil Science). The liners were cut with a circular saw and the sediment was split using a thin wire. After opening, the core halves were cleaned in a direction perpendicular to the core axis, to avoid contamination. The cores were logged with special attention to coarse layers, stones and deformation. One half was chosen to be photographed using a Canon EOS 400D and 15 cm sections of the core were photographed. These 15 cm images were stitched using the Corel PHOTO-PAINT software to get an image of the entire core. To improve the visibility, these images were noise-filtered and a colour- histogram equalization was applied with the Corel PHOTO-PAINT software. This is a method to increase the colour contrast of the image, by spreading the most frequent intensity values.

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Although this produces unrealistic colours, it is scientifically useful because it reveals layers one could miss on the real colour image. After photographing, the cores were packed, labeled and stored in a refrigerator at 4°C to prevent molding. One half core is treated as the work half, the other is the archive half.

7.4 Non-destructive core analyses

7.4.1 MSCL analyses Before opening the cores, they had been logged with a Geotek ltd. Multi-Sensor Core Logger (MSCL) at ETH Zurich, Switzerland, by P. Kempf. An MSCL is a fully automated device that comprises a rail on which the core is placed, and several non-destructive measuring sensors that can be mounted to this setup (Fig. 7.3). The MSCL of ETH Zurich consists of a gamma-ray attenuation density sensor, p-wave velocity sensor and a magnetic susceptibility loop sensor for full cores. The distance between the sensors and the set zero is precisely known, therefore the machine is able to correlate the points of measurements to the same depth points. A core pusher transports the core past the sensors in preset increments. The sampling increment used during the whole-core logging was 5 mm.

Once the cores were split, they were measured again with a magnetic susceptibility point sensor and a colour spectrophotometer at a 2 mm increment, this time at the Ledeganck campus of Ghent University (Department of Biology).

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Figure 7.3 This sketch of a Geotek ltd. Multi-Sensor Core Logger shows the central rail and some of the possible analysis devices (Geotek, 2012). Note that not all devices were used in this thesis.

7.4.1.1 Magnetic susceptibility Magnetic susceptibility is the bulk response of the sediment to the change of an external magnetic field. The potential of reinforcing or weakening the magnetic field is indicative for the type of material. For example, iron-titanium oxides and iron sulfides are ferromagnetic minerals with a positive magnetic susceptibility. Diamagnetic materials have a negative magnetic susceptibility. Examples are quartz and organic matter (Butler, 1992). The measurements on closed cores were done by pushing the core through a Bartington loop sensor (MS2C), mounted on the Geotek MSCL. The magnetic field is a low-intensity, non- saturating, alternating field (0.565 kHz). Split-core measurements were performed with a Bartington point sensor (MS2E). The sensor produces the same low-intensity, non-saturating, alternating magnetic field as the MS2C loop sensor, but with a strength of 2 kHz. The point sensor requires direct contact with the sediment, and can therefore only be used on split cores (Geotek, 2012).

7.4.1.2 Colour spectrophotometry Colour spectrophotometry describes the colour of a sediment sample in an objective and comprehensive way and provides an alternative to the more subjective Munsell colour system. It measures the reflectance for wavelengths between 360 and 740 nm, which corresponds to the very-near UV, visible and very-near IR range. A Konica Minolta CM-2600d

58 colour spectrophotometer, which produces a high-energy xenon flash illumination, is mounted on the MSCL. The measurement was done with a Ø 3 mm SAV aperture (Geotek, 2012).

7.4.1.3. Gamma-ray attenuation density Gamma-ray attenuation density is a method to measure the density of a core indirectly. A radioactive 137Cs source emits a narrow beam of gamma rays with energies at 0.662 MeV through an opening just above the sample guide rail. A detector is mounted on the other side of the rail to count the photons with the same principal energy as the source, thus only the transmitted photons that have not been scattered. The physical principle behind this method is Compton scattering, which is the primary mechanism of attenuation. Whenever a photon hits an electron it causes partial energy loss. The energy loss is directly proportional to the number of electrons in the core segment. Electrons have a negligible mass, but they occur in approximately equal amounts as protons. By knowing the electron density and the diameter of the core it is possible to determine the bulk density of the core material. (Geotek, 2012)

7.5 Destructive core analyses

7.5.1 Smear slides Smear slides are a cheap and fast way to study the composition of soft sediments. In each core, a number of layers were chosen, based on the histogram equalized images, so as to represent all the sedimentological facies and types of layers present in the core. These layers were studied in all cores to see changes in lateral distribution. A glass slide was labeled, and a drop of deionized water placed on this slide. On the predetermined location, a small amount of sediment, approx. 1 mm³, was extracted with the tip of a toothpick. The sediment was placed in the water drop and equally distributed over the slide. The amount of sediment needed is proportional to the grain size. A fine silt requires only a small sample (and more water), whereas a coarse sand requires a larger sample to achieve an ideal distribution. The slide is placed on a hot plate (ca. 60 °C) for a couple of minutes to let the water evaporate. When the slide is dry, a drop of optical adhesive is placed on the sediment. In this case the used optical adhesive was Canada balsam. Canada balsam is a synthetic resin often used for making microscope slides, as it has a refractive index similar to glass (n = 1.55) and its optical properties do not deteriorate with time. A cover glass is placed upon the slide and the balsam is gently distributed over the sediment. The best way to harden Canada balsam is by UV light. So the smear slides were placed in sunlight for a couple of hours. Excess balsam was excised with a cutter knife, and the slides were cleaned with acetone to dissolve the rest of the balsam. Since acetone leaves dirty evaporation traces, the glass was cleaned with ethanol afterward (Marsaglia et al., 2012).

The smear slides where studied under a normal optical microscope (Euromex) with 40x, 100x and 400x magnification. Smear slides have the disadvantage that the thickness of the grains cannot be flattened to a uniform thickness, as it is the case for thin sections. The variable

59 thickness causes a difference in birefringence within grains. To get a valuable description a procedure was followed to guarantee a coverage of the whole slide. Steps of 1 mm in the width and 2 mm in the length were taken, and what is seen in the center of the field of view was described. This prevents overestimation of the coarse fraction, as the describers attention is naturally attracted to the larger grains. It also guarantees a coverage of the entire slide, which is important since the distribution of the sediment on the glass is sometimes disproportional. In total an average of 200 picks were studied per smear slide.

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8. Results

8.1 Lake-floor morphology: Observations and interpretation The sonograms obtained during the sidescan-sonar survey have been combined into a mosaic to create a reflectivity map of the lake floor (Fig. 8.1). Some lines are of poor quality and had to be deleted, especially around the edges of the lake where the zodiac made a turn. Important structures are shown in more detail (Fig. 8.3). The reflectivity map has been modified into an interpretation map to show the most important features, (Fig. 8.5). The bottom track signal was used to calculate the depth below the tow fish, which is a measure for the water depth as the tow fish was towed at the water surface. These values were interpolated to create a bathymetry map (Fig. 8.2).

The first observation is that the size of the lake during the second field campaign is much smaller than on satellite images. This is not a seasonal change, because on all satellite images in Google Earth (taken on 08/02/2003, 29/04/2007, 03/01/2010 and 20/02/2011) the lake surface is larger. These images were all taken during the austral summer or autumn, the same season as that in which the sidescan-sonar survey was performed. The bathymetry in Salvetti (2006) shows depth values that are at least 2 m deeper than in the bathymetry in this study (Fig. 8.2).

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Figure 8.1 Reflectivity map of Lago Lo Encañado. The parts of the map in a red box are shown in more detail in Fig. 8.3.

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Figure 8.2 bathymetry map of Lago Lo Encañado. a) bathymetry map based on sidescan sonar data (this study). The core locations are included. b) bathymetry map based on depth soundings throughout the lake (adapted from Salvetti, 2006).

The highest reflectivity (light colours) occurs in the northern part of the lake, and along the eastern, western and southern slopes of the lake (Fig. 8.1). The darkest colours (low reflectivity) are found in the south-central part of the lake, which corresponds to the deepest part of the lake (Fig. 8.1; Fig 8.2). The straight white lines, most of them in a North-South direction, are the tracks along which the sonar fish was towed. The area along the shoreline shows high reflectivity values that have a spotted appearance. These spotted, high-reflectivity zones are interpreted as being caused by water plants (Fig. 8.3d). A photo from the eastern lake shore shows these water plants (Fig. 8.4).

The northern part can be separated into two areas. The northernmost area is shallow (white colours) (Fig. 8.1) and terminates with a steep slope between approx. 14 and 22 m depth, just offshore of the mouth of the Río Lo Encañado. South of this steep slope the lake is deeper and the slope is more gentle. The northernmost unit on the northern shore is the delta plain, the steep slope is the delta front, and below the steep slope is the prodelta (Postma, 1990).

From the northern shore, in an area where a fan-like feature occurs between approximately 10 and 20 m depth, at least four bright white linear features can be distinguished (Fig. 8.3c). Coinciding with the white linear features, the bathymetry contour lines show a positive morphology between 10 and 14 m depth and a negative morphology between 16 and 20 m depth, where the slope is steeper. This is again followed by a small positive morphology below the steepest part of the slope (22 m depth). At least one similar, but less pronounced, bright white linear feature occurs in the northeastern corner. Also in the southwestern corner of the lake similar structures occur. These high-reflectivity feature are interpreted as channels that coincide with the main river inflows. On the delta plain of the northern shore the sediment is

63 deposited from the river and forms a positive relief. On the steep delta front the subaqueous river is erosive and forms a channel. On the prodelta the sediment from the river forms a fan. Two small rivers enter the lake in the southwestern corner. They produce similar erosive channels as on the northern shore.

On several locations along the eastern and western shore (on the western shore and the northeastern, southeastern and southwestern corner), lobes of a high reflectivity occur on the slopes of the lake. These lobes produce a positive relief (Fig. 8.3b). A closer look shows that these lobes contain bright spots and concentric structures that look like ripples. The bright spots occur closer to the shore, whereas the ripples are seen more towards the center of the lake. The lobes are interpreted as a subaqueous continuation of colluvial fans that can be seen above water on satellite images (Fig. 8.1; Fig. 8.5). The ripples are interpreted as pressure ridges, positive and parallel structures that are the result of compressional stress on sediment fans (Bull et al., 2009). The bright spots on the fans are interpreted as large blocks.

Another high-reflectivity feature occurs along both the eastern and the western shore. A bright line runs parallel to the shoreline, but terminates when it crosses a bottom track. These high-reflectivity lines are no geological feature, but are artifacts. It is comparable to a multiple that can occur in seismic profiles. When a ping is emitted, only a small part is reflected towards the receiver. The rest of the signal is refracted away from the receiver, and is normally lost. In this situation, however, the refracted part of the signal hits the steep highly reflective shores and is reflected back towards the receiver. This signal arrives at the same time as the reflected part of the next ping. The result is a strong reflection that is received at the same time that the following ping should be received. The facts that the lines are exactly parallel to the coastline and that they terminate when they cross the bottom track support this interpretation.

One last feature is a structure in the center of the lake. It is characterized by discontinuous, sometimes bended bright lines, overall arranged in a concentric manner. This structure has a positive relief of about 2 m compared to the surrounding lake floor (Fig. 8.2; Fig 8.3a). It is elongated in a Northeast-Southwestern direction with an estimated length of 100 m and width of 50 m. At least the sides of this structure are covered with sediment (ENC07 is taken on that location and has a length of 55.3 cm). This large structure is difficult to explain. It could be an outcrop of bedrock. However, it is not likely to find such a structure in a valley that was cut out by a glacier. A more logical explanation is that this is a large block that slid from the steep slopes during a massive landslide. The location of the block is almost 400 m from each shore, which means that the block must have slid before the lake was filled with sediment. That places the timing of the landslide just after the deglaciation of that part of the valley. This phenomenon of slope instability after deglaciation has been observed in valleys in Austrian and French Alps (Kellerer-Pirklbauer & Kaufmann, 2007; Cossart et al., 2008). A seismic profile would be required to solve this problem and corroborate this interpretation.

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b)

65 c)

66 d)

Figure 8.3 Details of the reflectivity map (red boxes in Fig. 8.1). a) The structure in the center of the lake, viewed from two angles. b) A colluvial fan from the western shore, viewed from two different angles. In the left picture, the bright spots and ripples can be seen. In the right picture, a high reflectivity line can be seen which terminates as it crosses the bottom track. c) Three white lines in the center of the picture and one on the right side can be seen. The picture is taken at the northern shore. d) The high reflectivity line. Note that it runs parallel to the coastline and it terminates when it reaches the bottom track. The spotted, high-reflectivity zone along the coastline is created by plants and/or algae.

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Figure 8.4 View from the eastern shore on Lago Lo Encañado. The green line along the shore consists of water plants and/or algae. (Photo by P. Kempf).

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Figure 8.5 Interpretation of the sidescan-sonar map in Fig. 8.1.

8.2 Sediment cores: observations, analysis and interpretation In this chapter, the sediment cores that were used in this study will be described in detail and discussed. All sediment cores are characterized by a background sediment that is interrupted by a series of characteristic layers, which are interpreted as event layers (i.e. caused by processes or events that are different than those that cause the normal, background sediment). Four types of event layers can be identified in the cores: i.e. fining-upwards layers, white layers, blue layers and red layers. In the following paragraphs the main characteristics of the background sediment and of these event layers will be outlined to facilitate the description of the cores. The event layers will first be described on a macroscopic scale, based on visual identification and data from non-destructive analysis, and subsequently on a microscopic scale, based on smear slide analysis. Smear slides were taken on the most important locations, and studied cf. Marsaglia et al. (2012) and Myrbo et al. (2011). The term “fine fraction” will be used for everything that is visible with a 400x magnification, but is too small to be identifiable. The description of each layer will include information about the fine fraction.

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8.2.1 Background sediment

Macroscopic scale

The background sediment consists of alternating brown-green and white layers. The brown background sediment defines the average colour of the core. Sporadically the brown-green layer is replaced by a red or blue layer. The background sediment undergoes a transition from a strong blue colour at the base of the cores in comparison to the brown-green background sediment towards the top. The average level of MS and spectrophotometry data are higher in the lower (blue) part of the background sediment.

Microscopic scale

Most of the background sediment is made up of fine fraction, which consists of small grains and flakes of phyllosilicates. Most of the large grains are clods of organic matter and hematite. Besides these are minerals of quartz and feldspar. The darker layers contain more organic matter and clinopyroxenes. Diatom tests are intact and common. Like other layers the background sediment contains a small quantities of volcanic glass.

The difference between the brown-greenish background sediment in the upper part of the core and the bluish background sediment in the deeper parts of the cores has been visually described in the section above. The microscopic difference lies in the smaller quantities of organic matter, hematite, and clinopyroxene in the “bluish background”, which also contains more quartz. The fine fraction contains a large amount of colourless grains with a 1st order blue to grey birefringence colour. This could also be quartz, but the grains are too small to verify this. The difference between the upper (brown-greenish) and lower (bluish) sediment, and between different layers within background sediment is much smaller than expected based on the macroscopic description.

8.2.2 Fining-upwards layers

Macroscopic scale

This kind of layer consists of coarse-grained sediments and is characterized by a fining- upwards trend. The total thickness is variable, ranging from several millimetres to more than twenty centimetres. Its colour is grey to brown and dark grey to black after histogram equalization. Some fining-upwards layers have a blue colour. The grain size is coarse at the base, typically starting with an erosional surface. The fining-upwards trend includes a colour change to lighter brown. Sometimes a white layer caps the fining-upwards layer. There is no sharp boundary at the top, instead there is a gradual transition to the white layer. The fining- upwards layers in cores taken in proximity to the delta have a much coarser grained base than those in the more distal cores. In ENC01, the grains at the base of a fining-upwards layer are too large for a proper smear slide. Fining-upwards layers produce a positive MS peak. The

70 peak is steep at the base, and decreases gradually towards the top. On spectrophotometric lightness data these layers are characterized by a negative peak.

Microscopic scale

The fining-upwards trend that has been described in the macroscopic analysis is also identified in all smear slides. Most coarse layers have been sampled from base to top with three samples, so that a vertical fining-upwards trend can be confirmed. The grain size increases with proximity to the delta. The maximum grain size in the first fining-upwards layer in ENC01, on the delta slope, is 1 mm (ENC01; smear slide at 15.2 cm), the base of the first fining- upwards layer in ENC02 has a maximum grain size of approx. 0,2 mm. The top contains a large amount of fine-grained sediment. The top of the first major coarse layer in each core, except for ENC01, contains a large amount of diatoms, of which a small part, 5-10%, is broken. The amount of diatoms is highest at the top, and decreases towards the base, although they remain present. The coarse fraction consists mainly of quartz, feldspar (of which about half is plagioclase) and clinopyroxenes. The quartz minerals are colourless, but often overgrown by dark brown minerals, and the refractivity increases with grain size. Small grains have a white to blue colour (1st order) under crossed polarisers, but the refractivity increases to purple/orange (1st order) in larger grains. Feldspars are also colourless with low birefringence colours. They are often characterized by a “dirty” rim with a different birefringence. Most minerals are partly rounded. Euhedral crystals are very rare. The clinopyroxenes are green rounded crystals with a moderate relief and a low birefringence. The grain size increases towards the base of the coarse layer, and so does the number of types of minerals. Quartz, feldspars and clinopyroxenes are still present. The amount and size of opaque minerals increases strongly towards the base. Those minerals are often rounded, but few have a triangular shape or are associated with other minerals. The rims often have a brown red colour. The opaque minerals are hematite according to a description in Pichler & Schmitt- Riegraff (1997). Clods of brown organic materials (variable refracticity) and microlites (a colourless entity built up of smaller microcrystals with a mottled appearance under crossed polarisers) are present throughout the coarse layer.

Some phyllosilicates are large enough to be identified. There are green and white flakes (low relief, no birefringence) that have been identified as chlorite, muscovite and a smaller fraction of biotite. These minerals resemble the flakes that can be seen in the fine fraction. They are probably just larger flakes of the same material. Finally, the coarse layers contain plant remains (more in the proximal locations) and some volcanic glass. These can be identified as sharp concave-edged colourless grains that are amorphous. The amount of volcanic glass is very small.

The difference between the distal and proximal setting lies in the grain size. There is not too much difference in content. There are fewer phyllosilicates, more quartz and more plant material in the proximal (e.g. ENC01, on the delta) compared to distal (e.g. ENC02; in the centre of the lake) locations.

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8.2.3 Blue layers

Macroscopic scale

These layers have a distinct blue colour after manipulating the images with a colour histogram equalization. Its real colour is bluish grey, and not always easy to identify. The thickness is never more than 0.5 cm. The thickness is not linked with the distance of the core to the delta. The lower contact is sharper than the upper contact and not erosional. Sporadically the contacts are transitional. These blue layers cause a peak in spectrophotometry and they can cause a peak in MS data, but this is not always the case.

Microscopic scale

This type of layer consists mainly of fine fraction. A large part of this fine material is probably muscovite and chlorite, as these minerals also appear as larger flakes. The fine fraction also contains small pinkish red, amorphous grains with a high relief and sharp edges. This is volcanic glass. Diatoms are common, and so are clods of organic material and microlites. The large grains consist mainly of opaque minerals (hematite), but also feldspars, some quartz crystals and clinopyroxenes. Additionally there are shards of volcanic glass and plant remains in this larger grain-size fraction. The grain size increases towards the delta, and so does the abundance of clinopyroxene.

8.2.4 Red layers

Macroscopic scale

The red layers have a brownish red colour, and a bright red colour after colour histogram equalization. The darkest part is the centre. The top and bottom of this layer are diluted. The bright red layers show a small peak in MS, and a negative peak in spectrophotometric lightness.

Microscopic scale

The content of the red layers is similar to that of the blue layers. The main part is the fine fraction, small flakes of muscovite, chlorite and biotite. The biotite minerals are clearly pleochroic from ruby red to orange. These minerals are also present in larger grains. The larger grains are hematite, organic material, microlites, quartz, feldspars and clinopyroxenes. Diatoms are common, volcanic glass is present in small quantities. The red layers in cores close to the delta contain more and larger grains, and more clinopyroxene. The amount of biotite and hematite is much higher than in the blue layers, explaining the red colour. The opposite is true for clinopyroxene.

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8.2.5 White layers

Macroscopic scale

These layers are silty and have a greyish colour, which is white after colour histogram equalization. These white layers are deposited as part of the background sediment, but also occur in thicker layers above many fining-upwards layers. Because of this possible genetic link with the fining-upwards layers, the white layers are discussed individually from the background sediment. The thickness of the white layers above a fining-upwards layer is variable, ranging from 0.1 cm to more than 1 cm. The thickness is largest in the central park of the lake, both in absolute and relative (compared to the thickness of the coarse layer) thickness. The magnetic susceptibility data show a peak of low (but still positive) values, whereas the spectrophotometric lightness data show a positive peak.

Microscopic scale

These white layers are very fine grained. Most of the sediment consists of small transparent to green flakes that are probably chlorite and/or muscovite. The remainder of the sediment consists mainly of small dark and transparent grains, which are probably hematite and quartz or feldspar. Most of the larger grains are quartz and feldspars. There are small variations in grain size, but it appears that these changes are random across the lake. The changes in grain size are not related to the distance to the delta. The sediment contains a small amount of volcanic glass shards.

8.3 Core descriptions In this chapter each sediment core is described individually. The presence, thickness and shape of fining-upwards layers receive special attention, as do associated layers and deformed sedimentary units. The core images, core logs and MS, spectrophotometry and density data can be found in attachment (Fig 12.1-12.8). The given length of each core corresponds to the amount of the amount of sediment, excluding the oasis (the green foam) at the top of the core.

Core ENC01

This core was taken on the prodelta. It has a length of 63 cm. The top of the first fining- upwards layer occurs at 1 cm depth. This layer has a thickness of 23 cm. Above this fining- upwards layer are two white layers (each 0.2 cm thin) separated by a 0.1 cm thin brown layer. The base of this layer contains grains of fine gravel. The next fining-upwards layer is situated between 25.5 and 34 cm depth and has a 0.2 cm thin white layer on top. The sediment between these two fining-upwards layers is undisturbed. The sediment below the second fining-upwards layer, however, is gently folded and the next fining-upwards layer is situated

73 between 41.5 and 42 cm. It has a blue colour. Another fining-upwards layer is between 45 and 46.5 cm depth. It is folded and so is background sediment underneath. The lowermost fining- upwards layer is situated between 61.5 and 63 cm and has a blue colour. The sediment between the two lowest fining-upwards layers is heavily deformed and contains several sand lenses (Table 8.1).

Table 8.1: A recapitulation of the fining-upwards layers in ENC01 and ENC02. X = not present, V = very thin white layer (< 0.3 cm), VV = thin white layer (between 0.3 and 0.6 mm), VVV = thick white layer (>0.6 cm).

Top Base White layer Core ID (depth, in cm) (depth, in cm) on top? Comments ENC01 1 24 V gravel at the base ENC01 25.5 34 V ENC01 41.5 42 V blue colour ENC01 45 46.5 X folded ENC01 61,5 63 VV blue colour ENC02 1 4 X ENC02 7.3 8 X ENC02 14.5 16.5 VVV blue layer underneath ENC02 20.4 22.2 V white layer between two ENC02 24 27.9 X/V fining-upwards layers ENC02 38 38.5 VV blue colour ENC02 41 43.5 X ENC02 50 52 VVV blue colour ENC02 66 67.5 X combination of several ENC02 70 72.5 X coarse and white layers ENC02 77.5 80.8 V

Core ENC02

This is a 87 cm long core taken in the centre of the lake. It contains at least 10 fining-upwards layers. The uppermost fining-upwards layer is situated between 1 and 4 cm depth. The sediment below is weakly folded at the side of the core, but the laminae are clearly visible. There is a 0.7 cm thick fining-upwards layer at 8 cm depth. The third fining-upwards layer is 2 cm thick, with a 1 cm thick white layer on top and a 0.1 cm thin blue layer approx. 1 cm underneath. The next fining-upwards layer downwards at 22 cm depth also has a 0.3 cm thin white layer on top. The following fining-upwards layer contains a coarse-grained base, which is fining up towards a 0.3 cm thick white layer situated between 25.9 and 26.2 cm, above which the sediment coarsens again. There is no background sediment between the white layer and the fining-upwards layer on top. The next fining-upwards layer is situated at 38.5. There is no sharp termination at the top, but instead there is a gradual change from dark, coarse sediment to a white top, giving the whole unit a blue hue. At 43.5 cm is the base of the next fining-upwards layer. Another fining-upwards layer with dark coarse-grained base that gets diluted towards the top is situated between 50 and 52 cm. The white layer on top has a

74 thickness of 0.8 cm. There are two more fining-upwards layers at 66 - 67.5 and 70 - 72.5 cm depth. The latter contains at least three fining-upwards layers and two white layers that are intercalated. The last fining-upwards layer is situated between 77.5 and 80.8 cm depth, with a 0.3 cm thick white layer. The colour of the background sediment transitions from a brownish green to a blue hue below 49 cm. This core is considered to be the most complete record in this study. It shows no signs of disruptions or erosion. It will be used as the central core to which the other cores are correlated (Table 8.1).

Core ENC03

This core has a length of 84.3 cm. The core was taken in the deepest part of the lake. It contains at least 9 fining-upwards layers. The first fining-upwards layer of 2.2 cm thickness is found between 0.8 and 3 cm depth, followed by a 3.5 cm thick coarse layer (4.5-8 cm depth). The third fining-upwards layer is characterized by a 1 cm thick white layer on top of the fining- upwards layer (18.2-20 cm), and a 0.3 cm thin blue layer at approx. 1 cm below the fining- upwards layer. Between 24.5 and 26 cm depth is the next fining-upwards layer. A unit composed of at least two fining-upwards layers with an intercalated white layer is situated between 30 and 32.5 cm. The next fining-upwards layer is found between 41 and 42 cm. The base is coarse-grained and dark in colour. The colour changes gradually to blue and further to white towards the top. A coarse layer of 2.2 cm thickness is situated between 44.3 and 46.5 cm. The next fining-upwards layer (53.5-55 cm) has a blue colour and a 0.6 cm thick white layer on top. The last fining-upwards layer is found from 80.2 cm until the bottom of the core (84.3 cm). This last fining-upwards layer has a 0.2 cm thin white layer on top. There are no recognisably homogenous layers or deformed layers. The background is finely laminated with alternated white and coloured layers (red, blue, brown). Under the fining-upwards layer at 46.5 cm the background colour becomes bluer as there are more blue layers and thicker white layers. At least 5 clear red layers can be identified (Table 8.2).

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Table 8.2: A recapitulation of the fining-upwards layers in ENC03 and ENC05. X = not present, V = very thin white layer (< 0.3 cm), VV = thin white layer (between 0.3 and 0.6 mm), VVV = thick white layer (>0.6 cm).

Top Base White layer core ID (depth, in cm) (depth, in cm) on top? Comments ENC03 0.8 3 X ENC03 4.5 8 X ENC03 18.2 20 VVV blue layer underneath ENC03 24.5 26 X ENC03 30 32.5 X white layer intercalated ENC03 41 42 V blue colour ENC03 44.3 46.5 V ENC03 53.5 55 VVV blue colour ENC03 80.2 84.3 VVV ENC05 0.5 1.8 X ENC05 3.5 5.5 V ENC05 32 34.4 VVV blue layer underneath ENC05 38.5 43 X

Core ENC05

This 43.1 cm core was taken in the central western part of the lake. The first fining-upwards layer has a thickness of 1.3 cm thick fining-upwards layer with a base at 1.8 cm depth. The sediment below is laminated. The next fining-upwards layer (3.5-5.5 cm) has a thin white layer on top. The next 26.5 cm of background sediment below this coarse layer are deformed and folded. This unit contains some sand lenses. The fining-upwards layer below (32-34.4 cm) has a white layer (0.4 cm thick) on top and a blue layer (0.2 cm thick) below. The sediment below this coarse layer is laminated. The lowermost fining-upwards layer (38.5-43 cm) contains a large mud clast (> 1 cm diameter). This core does not contain any blue coloured fining- upwards layers (Table 8.2).

Core ENC06

This is the westernmost core taken on the delta slope. It has a length of 68.4 cm. The first fining-upwards layer of approx. 1 cm is weakly folded, and so is the silty sediment below. The 6 cm thick fining-upwards layer below is horizontal at the bottom, and folded at the top. The sediment below is homogenized or deformed until a 0.4 cm thick fining-upwards layer at 54.6 cm depth. The deformed sediment contains at least one fining-upwards layer (29-31 cm) that holds a rounded pebble, and several sand lenses. The sediment below the deformed unit is weakly folded and is clearly laminated. The next fining-upwards layer has an erosional base (56.3 -58 cm) and no white layer on top. The lowermost fining-upwards layer is blue in colour

76 and has a white layer on top (62.5– 65 cm). There is a sub-angular pebble grain of about 2 cm diameter between the two fining-upwards layers (Table 8.3).

Table 8.3: A recapitulation of the fining-upwards layers in ENC06 and ENC07. Note that fining-upwards layers in deformed units are not included. X = not present, V = very thin white layer (< 0.3 cm), VV = thin white layer (between 0.3 and 0.6 mm), VVV = thick white layer (>0.6 cm).

Core ID top (depth, in cm) base (depth, in cm) white layer on top? comments ENC06 1.5 2.5 X folded ENC06 4 10 X folded top ENC06 54.2 54.6 X blue colour ENC06 56.3 58 X ENC06 62.5 65 VVV blue colour ENC07 0.2 0.5 X ENC07 2.5 5.5 V

Core ENC07

The core was taken on the prodelta on the block in the center of the lake (Fig 8.3a). The length of the core is 55.3 cm. The core starts with a 0.3 cm thin fining-upwards layer at 0.5 cm depth. The next 3 cm thick fining-upwards layer (2.5-5.5 cm) has a 0.1 cm thin white layer on top. The remainder of the core is deformed or homogenized. In the middle of this deformed unit is another fining-upwards layer. It is folded and holds a purple very coarse gravel grain of about 4 cm diameter, probably of andesite (Table 8.3).

Core ENC08

This core was recovered from the north-western part of the lake and contains 72.8 cm of sediment. A folded fining-upwards layer is seen at 2.2 cm depth. It does not have a white layer on top and therefore it is difficult to see the upper limit of this layer. It is followed by 6 cm thick fining-upwards layer (4.5-10.5 cm). This layer consists of at least two, possibly three relative peaks in maximum grain size with low MS layers intercalated. On top is a thin white layer (0.2 cm thickness). The 10 cm of sediment below are deformed. This deformation extends to a fining-upwards layer at 24 to 24.5 cm depth, with a white layer on top with approx. the same thickness. The next fining-upwards layer is between 28 and 30 cm depth. There are three more fining-upwards layers at 44.5-44.8; 62-63.5 and 64-67 cm. The 62-63.5 cm layer might consist of two different fining-upwards layers separated by a thin, white layer. The last fining-upwards layer (64-67 cm) is folded with a 0.8 cm thick white layer on top. Below 57 cm the background sediment is blue. There are blue coloured fining-upwards layers in this core (Table 8.4).

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Table 8.4: A recapitulation of the fining-upwards layers in ENC08 and ENC09. Fining-upwards layers in deformed units are not included. X = not present, V = very thin white layer (< 0.3 cm), VV = thin white layer (between 0.3 and 0.6 mm), VVV = thick white layer (>0.6 cm).

Top Base White layer on Core ID (depth, in cm) (depth, in cm) top? Comments ENC08 0 ? 2.2 X 3 fining-upwards trends within ENC08 4.5 10.5 V fining-upwards layer ENC08 24 24.5 VVV ENC08 28 30 X dark layer underneath ENC08 44.5 44.8 X two coarse layers separated by ENC08 62 63.5 X white layer ENC08 64 67 VVV folded ENC09 0.5 1 X ENC09 3.5 7.5 V ENC09 28 31 VVV blue layer underneath ENC09 34.5 36.5 V ENC09 39 42 V ENC09 50 50.8 V ENC09 65.5 67 X blue colour ENC09 83.5 85 X ENC09 89 89,8 ENC09 91.8 94.5 V

Core ENC09

This is the south-westernmost and longest core of this study. It has a length of 102.1 cm. At 1 cm depth there is a 0.5 cm thin fining-upwards layer. The next fining-upwards layer (3.5- 7.5 cm) has a 0.2 cm thin white layer on top. The 20 cm of sediment below that fining-upwards layer are deformed. In this deformed unit is a folded sequence of (from base to top) a blue layer, a coarse layer and a white layer. This sequence (from base to top blue-coarse-white) with similar thickness, yet undeformed, lies below the deformed unit (28-31 cm). The third coarse layer thus seems to be a repetition of the fourth coarse layer. There is another fining- upwards layer between 34.5 and 36.5 cm, followed by two fining-upwards layers separated by a white layer between 39 and 42 cm. Between 50 and 50.8 cm is the next fining-upwards layer, with a 0.1 cm thin white layer on top. The sediment below this fining-upwards layer is deformed. At a depth between 64.8 and 67 cm, there is a blue fining-upwards layer with a 0.5 cm white layer on top. The next 10 cm of sediment below are deformed. There are three more coarse layers, at 85, 89.9 and 94.5 cm. Only the latter has a 0.2 cm thin white layer on top, and an erosional base. The sediment below this last fining-upwards layer is deformed. The lower part of the core has a rubiginous appearance (Table 8.4).

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9. Interpretation & Discussion

9.1 Interpretation of the different layers Fining-upwards layers

The fining-upwards layers can be found and correlated throughout the lake. They contain high amounts of sand, and some pebbles. The large range, the large grain size at the base and the fining-upwards trend can only be explained by an (at least partly) subaqueous slide. The deposits are therefore turbidites in the loose definition of the term (chapter 2). These coarse layers will from this point on be referred to as turbidites. The uppermost turbidite in ENC01 is the largest turbidite deposit in any of the cores in this study. Comparing this to the Bouma sequence in Shanmugam (2000) shows that this turbidite contains the complete Bouma sequence (Fig. 9.1). Unit Ta is the largest unit in this turbidite. It is possible to see the fining- upwards trend. It is not possible to see the more detailed Stow & Shanmugam subdivisions in the upper part of the turbidite, and the Lowe units are not present. For comparison, the uppermost turbidite in ENC02 is also shown. In this core, the coarse part of the turbidite is condensed in comparison to that in ENC01.

Figure 9.1 A comparison between the Bouma sequence and the upper turbidite in ENC01 gives an excellent match. The lower part of the turbidite (Ta-Td) is condensed in the same event in ENC02. Adapted from Shanmugam (2000).

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More precisely, these turbidites can be classified as LT2 turbidites, according to the classification in Van Daele (2013), as they are composed of dark terrigenous material and thus differ from the background sediment. There are two possible sources for this terrigenous sediment. Either it comes from a collapse of part of the delta, which is fed with sediments from within the catchment area by the river. The other option is a landslide surging directly into the lake, preferably from the steep slopes to the east and west of the lake. A third option, that of a flood in the catchment carrying larger grains into the lake, can be dismissed because the large pebbles that can be found at the base of the turbidites cannot be transported to the prodelta by a flood. Furthermore the presence of MTD deposits below some turbidites can only be explained by a slump in the lake. As mentioned, a delta or slope collapse can be the result of an earthquake, but cannot be used as proof of an earthquake. There are however strong indications of an earthquake trigger in some of the larger turbidites. The fact that some coarse layers comprise multiple MS peaks (interpreted as multiple fining-upwards sequences) supports the hypothesis of a seismic triggering. It suggests that there were either multiple shocks in a very short time frame, or more likely that there were multiple turbidites triggered at the same time in different parts of the lake. Both can only be explained by an earthquake. Most turbidites are coarse and dark, but the relative thickness of the white layers that cover them is not constant.

Red layers

The red layers are interpreted as erosive material coming from the rocks in the proximity of the lake. On satellite photos there is a change in colour that coincides with a boundary on the geological map (Fig. 4.6). The red sediment comes from unit OM2C, which consists of volcanosedimentary deposits of basaltic, dacitic, epiclastic and pyroclastic rocks. Mulder et al. (2001) describe inversely graded turbidites in the Mediterranean sea. They link this to periods of higher discharge of rivers, defining these layers as flood turbidites. The base of the red layers is lighter than the central part, unlike the turbidites where the base is the darkest part and the colour becomes lighter towards the top. Therefore they are possibly inversely graded, and a similar process as the floods in Mulder et al. (2001) could be responsible for the creation of the red layers.

Blue layers

The blue layers are interpreted to have a similar origin as the red layers, but come from the more distant parts of the catchment area. They contain erosive material from the other side of the geological boundary, also seen in satellite images. This geological unit is M3i, which comprises partially eroded volcanic complexes together with andesitic, basaltic to dacitic lavas, breccias, domes and pyroclastic deposits. This unit lies for the greatest part in the catchment of Laguna Negra, the lake to the east of Lago Lo Encañado. A core from Laguna Negra, taken in the basin west of the peninsula, shows a clear blue colour (Fig. 9.2). In the microscopic description of the two layers in chapter 8.3 is shown that both layers have the same mineralogical content, but that the different quantities produce the difference in colour.

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This is consistent with the hypothesis of two separate origins within the same catchment area. The increase of sediment influx from erosive material in one part of the catchment does not prevent the influx of erosive material from the other part. The reason for these sudden changes in sediment influx is unclear. It could be weather induced. A period of higher melt water runoff from the glacier could increase the volume of “blue” sediment. Another explanation can be that a landslide somewhere in the catchment (triggered by an earthquake or not) produces a sudden increased input of clastic material of that type of rock.

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Figure 9.2 A comparison between a core taken from Laguna Negra (NEG02) and Lago Lo Encañado (ENC03). The light blue layers are more frequent in the core from Laguna Negra. The background colour in ENC03 turns blue below the turbidite at 46.5 cm. Both core images are adjusted with histogram equalisation.

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White layers

White layers are composed of very fine-grained sediments. These layers occur both within the background sediment and on top of large turbidites. They are probably deposits of hypopycnal flows. During periods of normal sedimentation they enter the lake through the river and are dispersed as buoyant plumes. When a turbidity current occurs, sediment is remobilized. The coarse fraction is transported as a hyperpycnal density flow along the lake floor, whereas the fine fraction forms a hypopycnal or homopycnal density flow above the former hyperpycnal current (Mulder & Alexander, 2001; Fig. 2.1). This is in agreement with the observation that the relative thickness of the white layer increases towards the center of the lake in relation to the thickness of the turbidite deposit.

Background sediment

The background sediment is a collective noun we have used to describe the laminated deposits between two event deposits. It contains red, blue, white and brownish green layers. These have been described individually above. The background sediment consists of a repetition of white layers followed by coloured layers. This repetition is presumably seasonal. The white layer is interpreted as having been deposited during austral winter, when the lake is frozen and the finest sediments are deposited. The coloured layers are deposited during austral summer when the lake is ice-free and sediment is washed into the lake by melt water from the catchment. They can therefore be interpreted as varves. Varve counting between two dated turbidites produced dates that are comparable to the age modeled dates (see further). They are however not very accurate, mainly due to a low resolution, maybe also due to erosion. For example, between the first two turbidites, 23 combinations of a white and a coloured layer can be counted, interpreted as a deposit of 23 years. In one of the following chapters (chapter 9.4.1) these turbidites will be dated at 2010 and 1985 respectively, which means there is a time span of 25 years between the two turbidites. One other important observation of the background sediment as a whole is that there is a rapid colour change from bluish white (bottom of the cores) to brownish green (top of the cores). This is caused by an increase in organic matter. At the same time the size of the individual layers decreased, which indicates a decrease in sedimentation rate. This could be explained by a retreating glacier and an increase of the average temperature in the valley. The valley and the lake itself have thus become a more gentile environment, which increased plant and algae growth in the glacial valley and on the delta. This has increased the amount of organic matter in the lake and therefore the browner colour of the sediment. A smaller glacier also produces less erosive material and soil formation in the catchment following the glacier retreat may have fixated easily erodible sediments, thus decreasing the sedimentation rate, especially from the blue sediment. This can however not be proven as the glacier has only been monitored since 1975 (Escobar et al., 1995). A link between the change in sediment colour and the construction of the dam can be excluded. The lake has been dammed since the 1950’s (Salvetti, 2006), and the change in colour occurred in the 19th century, as shown in chapter 9.4.2.

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9.2 Correlation of the turbidites A first identification and correlation of events was made visually, aided by MS and spectrophotometry data. The cores were also correlated with a core from Salvetti (2006), who dated the core with two 14C dates, and an additional 210Pb chronology on the upper part of the core.

The cores have been aligned into two, more or less West-East oriented cross-sections. The northern cross-section consists of cores taken on the delta, namely ENC06, ENC01, ENC07 and ENC08, from West to East. The southern cross-section contains ENC05, ENC02, ENC03 and ENC09. The core in Salvetti (2006), core 1305, is located between the two cross-sections. The sidescan sonar bathymetry suggests that the core is taken on the central part of the lake and not on the delta.

The southern cross-section

This series of cores is discussed first, as it gives a more complete event stratigraphy (Fig. 9.3). In cores ENC05-03-09 there is a turbidite at the top, which is missing is ENC02. It has a thin, coarse base, but is generally finer in grain size than other turbidites. This turbidite is labeled event A. The next turbidite (event B) is characterized by a thin white layer on top and a thickness of at least 3 cm. In ENC05 and ENC09 this turbidite is covering an MTD deposit. This can be recognized by a folded repetition of event E in ENC09. Event C can only be distinguished in ENC02 and ENC03 due to erosion by the MTDs of event B in ENC05 and ENC09. It is a dark layer with coarse base, it lacks a white layer on top. Below is another thick turbidite (event D) that can be seen in all four cores and also in core 1305. Typical for this turbidite is the thick white layer of about one cm on top, and a thin blue line underneath. This blue layer is not genetically related to the turbidite, but is mentioned because it provides extra support for the correlation between cores. The next turbidite, called event E, can also be identified in all cores by the thin white lamination just at the top of the turbidite. It is the lowermost turbidite in core ENC05, and the turbidite above the homogenite in core 1305. Events F and G can be seen in cores ENC02, 03 and 09 as two coarse layers divided by a white layer. It is unclear if the two events are related (two turbidites from the same earthquake or an earthquake followed by an aftershock) or are two small separate events that happened within a short period of time. The white layer between the two turbidites suggests that there was some time between the two events. The lower layer is the main turbidite, especially in ENC09, where it is has a coarse base and a thick white layer on top. The layer above is thin and not as coarse in grain size. It is possible that only the lower part was deposited by an earthquake-triggered slump directly in the lake, and the upper part is formed by onshore landslide deposits (from the same earthquake) that were washed into the lake at a later stage. Event H is again recognizable by a white layer on top of the turbidite and a dark layer underneath. This turbidite is blue in colour. The next event, labeled event I, is only clearly seen in ENC02 and 03, but can be recognized in ENC09 by a peak in the MS values. The same is true for event J. The turbidite is blue and has a thick white layer on top, and is recognized by a

84 high MS peak. Event K is represented by a coarse layer with at least two MS peaks in ENC09, and by at least two separate coarse layers divided by a white layer. Event L is also shown by two coarse layer and a white layer in between. Event M is the lowermost turbidite. It has a white layer on top and a thickness of at least 2 cm. Below this turbidite in ENC09 is another MTD.

Figure 9.3 A correlation of the turbidites in the cores on the southern cross-section. Core 1305 is taken from Salvetti (2006). A to M show the event deposits as described in the text.

The northern cross-section

The cores are characterized by erosive layers, thick deformed units and partly homogenized sections. The correlation is therefore a difficult process. Core ENC02 has been added to the

85 figure to make the correlation a little clearer. We intend to keep the same labeling of the turbidites as in the previous section (Fig. 9.4).

The uppermost event, event A, is present in cores ENC06-07-08, but not in ENC01-02. This layer is very coarse grained in ENC06, whereas is ENC07 this is only a minor deposit. This can probably be explained by the location of this core, on a high point in the bathymetry (Fig. 8.2). The second event (B) occurs in every core, although the thickness varies strongly. The base of this turbidite consists of gravel in ENC01, and coarse sand in ENC06 and ENC08. An MTD is situated below this turbidite in ENC07. Below this point the correlation is not as straightforward as for the southern line. The following correlations should therefore be taken with precaution. The second turbidite in ENC01 is identified as event E, because of the sharp white layer on top. It is very coarse-grained (gravel to coarse sand at the base). Events B and E were both highly erosive in ENC01, because events C, D, F and G have been eroded. Events H and I can be correlated between ENC01, ENC02 and ENC06, because of the blue colour of the turbidite, the white layer on top and a thin, dark layer below (H), and because of a coarse layer which lacks a clear white layer on top, but has two dark layers underneath (I). There is an MTD below the turbidite linked to event I in ENC01 and ENC06. Event J is recognized in ENC01 and ENC06 because of the blue colour.

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Figure 9.4 A correlation of the turbidites in the cores on the northern cross-section. ENC02 is added to help the correlation.

9.3 The intensity of an earthquake In this chapter we try to correlate the turbidites in our cores with recent earthquakes. The most important factor for turbidite formation is the local seismic-intensity value. As the study area is sparsely populated, historic records lack seismic-intensity values for the Maipo Valley. However, some authors have empirically developed an equation to estimate the intensity based on the magnitude and the distance to the epicenter or the rupture zone. For intra- continental-plate earthquakes, the MMI was calculated using the following equation from Bakun & Wentworth (1997):

MMI = 5.07 + 1.09Mw - 3.69Log10 (De)

The formula uses the Mw and the distance to the epicenter De of an earthquake. The dataset used to develop this formula contains shallow (<20 km) intra-continental-plate earthquakes in

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California with a Mw between 4.7 and 6.9. For earthquakes within this magnitude range, we expect this formula to be a good indication of the real intensity.

The MMI for interplate (i.e. subduction) earthquakes was calculated using the following equation, from Barrientos (1980; in Barrientos, 2007):

MMI = 1.3844Ms – 3.7355Log10 (Dr) – 0.0006Dr + 3.91

This formula is based on the Ms and the closest distance to the rupture zone (Dr). Seismic intensity is traditionally expressed in Roman numerals, but since these calculations will yield decimal numbers it was opted to use Arabic numerals instead.

There is no universally accepted threshold above which seismites are triggered. Van Daele (2013) on Chilean lakes, and Monecke et al. (2004) on Swiss lakes report intensity thresholds of VI and VI-VII, respectively. The data used to calculate intensities hold a high uncertainty, especially on the rupture zones of the older earthquakes. The intensity values presented in the following section are therefore estimates that help us appreciate the possibility of generating a turbidite as a result of this earthquake.

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Table 9.1 Summary of all known historical earthquakes in Central Chile between 2012 and 1575, together with the location, source in literature, type, magnitude, approx. distance to Lago Lo Encañado and estimated seismic intensity. The distance is the closest distance to the rupture zone (Dr) for subduction earthquakes or the distance to the epicenter (De) for intraplate earthquakes. Ms values are used for subduction earthquakes and Mw values are used for intraplate earthquakes. All earthquakes with an seismic intensity above 6 are highlighted.

Distance calculated MMI Year EQ Source Type Ms Mw (De) for LLE 2012 Maipo Valley USGS intraplate 3 4 15 5 2010 Maule Pulido et al., 2011; Wang et al., 2012 subduction 8.5 8.8 185 7.1 1985 Valparaíso Barrientos, 1988 subduction 7.8 175 6.2 1960 Valdivia Kanamori & Cipar, 1974; Lomnitz, 2004 subduction 8.3 9.5 520 4.9 1958 Maipo Valley Lomnitz, 1960; Alvarado et al., 2009 intraplate 6.9 6.3 18 7.3 1943 Illapel Lomnitz, 1970b; Beck et al., 1998 subduction 7.9 - 8.3 350 5.1 - 5.7 1939 Chillan Lomnitz, 1970b; Leyton et al., 2009 intraplate 7.8 - 8.3 7.8 - 8.3 340 4.2 - 4.8 1906 Valparaíso Okal et al., 2005 subduction 8.3 175 6.9 1861 Mendoza Lomnitz, 1970a; Perucca & Moreiras, 2006 intraplate 7.2 7.2 140 5 1850 Maipo Valley Lomnitz, 1970b intraplate 7 - 7.5 7 - 7.5 18 8.1 - 8.6 1835 Concepción Lomnitz, 2004 subduction 8 - 8.5 300 5.6 - 6.3 1829 Valparaíso Lomnitz, 2004 subduction 7 175 5.1 1822 Valparaíso Lomnitz, 2004 subduction 8 - 8.5 175 6.5 - 7.2 1751 Concepción Lomnitz, 2004 subduction 8.5 250 6.6 1730 Valparaíso Lomnitz, 2004; Udías et al., 2012 subduction 8.5 - 9 175 7.2 - 7.9 1657 Concepción Lomnitz, 2004; Udías et al., 2012 subduction 8 380 5.1 1647 Valparaíso Lomnitz, 1970b; Udías et al., 2012 subduction 8 - 8.5 175 6.5 - 7.2 1580 ? near Santiago Cisternas et al., 2012 ? ? > 6.2 1575 ? Near La Ligua Lomnitz, 1970b ? 7 - 7.5 140 4.8 - 6.2

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9.4 Historical earthquakes and their deposits

9.4.1. 20th and 21st century The 210Pb age model in Salvetti (2006) allows to date the turbidites in our cores and to compare them with the earthquake record of the last century. The record in the sediment archive matches very well with the historical evidence. We therefore conclude that these layers are in fact turbidites triggered by seismic activity.

Event A

Cores ENC02 and ENC03 are taken in each other’s proximity and are therefore similar. Yet ENC03 contains a small turbidite at the top of the core (also visible in ENC04-09; i.e. cores from the second field campaign), which is missing in ENC02 and ENC01 (i.e. cores from the first field campaign). Hence, this turbidite was formed in between the two field campaigns, more precisely between 15/01/2011 and 21/02/2012 (Fig. 9.5). Moreover, if this is an earthquake-triggered turbidite, this earthquake must have occurred in this time window. The most likely earthquake to have triggered the

turbidite was a local Mw 4.0 intraplate earthquake on the 13th of January 2012 with an epicenter approx. 15 km southeast of the lake and a depth of 14 km. The intensity on the MMI scale was calculated using the equation in Bakun & Wentworth (1997). This yields an intensity of 5.1 (Table 9.1), which is lower than the empirical intensity threshold of VI. An explanation might be the fact that there was an important drop in the lake level between the two surveys. This is also visible when comparing the contours of the lake on satellite pictures and the contours on the sidescan sonar data (Fig. 8.1). In this period the region of San José de Maipo was hit by a serious drought. Various local newspapers report that this region was declared to have a water deficit, that glaciers were melting and springs were drying up. In February 2012 the government decreed the region around San José de Maipo as a disaster area because of the drought. A disequilibrium can occur after a fast lake-level drop, creating instable lake sediments (Fanetti et al., 2008; Moernaut et al., 2010). In this context, a low- intensity earthquake could possibly create a turbidite.

Another explanation is that this is not a turbidite that ran directly into the lake, but an indirect influx of material from debris landslides in the catchment area triggered by the 2010 Maule earthquake. We doubt this explanation because there is almost one year between the 2010 earthquake and the first field campaign, and no sign of this influx in ENC01 and ENC02. Moreover, we would expect to see this type of deposits after every major earthquake, and that is not the case here (see further).

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Figure 9.5 A comparison between ENC03, ENC02 and 1305 (Salvetti, 2006). The year between brackets is the year in which the core was taken. This shows that event deposit A was formed between (January) 2011 and (February) 2012, and that the turbidite linked to event B was deposited between 2011 and 2005.

Event B

The next large turbidite in the sequence occurred between 2006 (no turbidite in core 1305) and January 2011 (first field campaign) (Fig. 9.5). In most cores this earthquake deposit is characterized by a thin white layer on top. In cores ENC05 and ENC09 this event deposit includes both a turbidite and a deformed unit below. Because of the size of the turbidite and the timing this is almost certainly the deposit caused by the 2010 Maule earthquake. This is in agreement with the interpretation of Van Daele (2013). In ENC08 the turbidite contains multiple fining-upwards sequences, which we interpret as multiple turbidity flows triggered by the same event, resulting in a stacked turbidite similar to the stacked turbidites of Van Daele et al. (in press). The rupture

distance is estimated at 185 km using a slip model in Pulido et al. (2011), and the Ms is 8.5. Van Daele (2013) calculated a Seismic Intensity of VI-VII for this earthquake, which is strong enough to trigger a turbidite. Using the equation in Barrientos (1980), the MMI is calculated at 7.1 (Table 9.1).

Event C

There is a small turbidite halfway between the 2010 and the 1958 earthquake (see event D). Interpolation yields an estimated age of 25 years before 2010. This is in agreement with the age-depth model in Salvetti (2006). The earthquake must therefore have happened around 1985. The best match in the USGS earthquake archive is an earthquake on the 3rd of March 1985 offshore Valparaiso. The rupture

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zone of this Mw 7.8 interplate earthquake was located at an estimated distance of 175 km from the lake. The surface-wave magnitude of this earthquake can be assumed at

Ms 7.8, as the moment magnitude and surface-wave magnitude are equal between 6.2 and 8.2 (Scordilis, 2006). The MMI is estimated at 6.2 (Table 9.1).

Event D

The next turbidite is characterized by a white layer above, and a thin blue layer below. There is a peak in 137Cs in core 1305 just above the turbidite, which coincides with a peak in radioactive fall-out. This peak corresponds with the Southern Hemisphere fall- out maximum between 1963 and 1965 (The Environmental Measurements Laboratory, 2008; in Von Gunten et al., 2009a). It is the result of extensive atmospheric nuclear weapon tests which reached a peak in 1962, until it was banned by an UN treaty in 1963 (United Nations Office of Disarmament Affairs).

The first option to link this turbidite to is the 1960 Valdivia earthquake (Mw 9.5; Ms 8.3). The distance between the rupture zone and the lake is estimated at 520 km. This yields a seismic intensity of 4.9. Van Daele (2013) used a seismic intensity of 5 after Lazo (2008), and saw no deposits from this earthquake. The 1960 Valdivia earthquake, despite being the largest ever recorded earthquake, probably ruptured too far to the south to produce a turbidite. A more likely trigger for this turbidite is the 1958 Maipo Valley or Las Melosas earthquake. The earthquake, which consisted of three separate

shocks, had an Ms between 6.7 and 6.9 (Flores, 1960). These magnitudes were

recalculated to an Mw 6.3 (Alvarado et al., 2009) The distance between the epicenter and Lago Lo Encañado is approx. 18 km. This yields an intensity of 7.3 on the MMI scale, which is in agreement with the MSK seismic intensity map in Alvarado et al. (2009).

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Figure 9.6 Map of the Maipo Valley showing the seismic intensity (MSK) values after the 1958 Maipo Valley earthquake. LLE = Lago Lo Encanado (enlarged for visibility purposes). The seismic intensity was between VII and VIII around the lake. Adapted from (Alvarado et al., 2009)

Event E

The sediment just above the turbidite and the homogenite in core 1305 has a dated age of 100 years before coring. As the core was taken in 2005 this event occurred at the beginning of the 20th century. The turbidite can be recognized in other cores because of its thin, but prominent, white layer on top. The most likely trigger for this earthquake was the 1906 Valparaíso earthquake. As explained in chapter 5.1.1., the uncertainty for this earthquake is extremely high because of the coincidence with an

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earthquake in the Aleutian Islands. We follow the estimation in Comte et al. (1986) that the location of the rupture zone was similar to the 1985 Valparaíso earthquake. The distance between the rupture and the lake is therefore also estimated at 175 km.

The surface-wave magnitude is Ms 8.3. This yields a calculated MMI of 6.9 (Table 9.1).

9.4.2. Comments on the 20th and 21st century earthquakes To our best knowledge, all major earthquakes in Central Chile that occurred between 2010 and 1906 and produced a large enough intensity (i.e. an intensity above VI on the MMI scale) can be identified in the sediment archive of Lago Lo Encañado.

9.4.3. Pre-1900 earthquakes From the previous paragraphs, we can conclude that the sediments of Lago Lo Encañado represent a sensitive archive of the seismicity during the past century. Below 1906, core 1305 shows a hiatus because of a homogenite. The 14C dates below this homogenite are much older than expected (1002 AD). This arouses suspicion as it would give a hiatus of 900 years. However, in cores ENC02 and ENC03 (central, deepest part of the lake) there are no indications of such a large hiatus. Therefore, we reject the 14C dates of Salvetti (2006) and attribute the old ages to a reservoir effect. This means that we lack an age model. However, we have no reason to believe that there are major changes in the sedimentation rate of the lake between event deposit E and event deposit I, the point below which the background sediment changes to a more blue colour. Therefore, an attempt to date older turbidites can be made by extrapolating the sedimentation rates of the last century.

Assuming that the intensity threshold of VI was also valid prior to 1906, there were at least 5 earthquakes that are expected to have caused a turbidite in the time period from 1550-1900 AD. For the earthquakes on the Valparaiso segment, a rupture distance of 175 km has been estimated. This is the same distance we used for the 1985 earthquake on the same segment. The uncertainty of the intensity estimations becomes larger with increasing age, because the uncertainty of the magnitude and the rupture zone increases.

In a first attempt, an “age-depth model” was created by using the average sediment rate in the past 104 years to calculate an age for the turbidites below the 1906 turbidite. Using this strategy did not result in the desired match of the turbidites and the historical archive. However, these poor results can be explained. First of all, the sedimentation rate in the upper 104 years is not constant. In ENC02, the average sedimentation rate in the center of the lake between 2010 and 1985 was 0.128 cm/year, between 1985 and 1958 it was 0.233 cm/year and between 1958 and 1906 it was reduced to 0.075 cm/year. However, also changing the extrapolated sedimentation rate to any sedimentation rate in this window does not result in a match between turbidites and historical earthquakes. A reason why these constant- sedimentation-rate methods do not give the desired results could be that the background sediment changes below event layer I. The reason for this change has been speculated upon in chapter 9.1. We believe that there is a higher sediment influx below event I. Events F and G were combined as one event to simplify the model. A “two sedimentation rates model”, with

94 a first sedimentation rate SR1 between events E (1906) and I, and with a second sedimentation rate SR2 below event I was developed. The sediment rate SR1 was empirically set at 0.167 cm/year, and the sediment rate SR2 was set at 0.294 cm/year. This is consistent with the assumption that the sediment rate increased below event I. With this age-depth model it is possible to link events to the historical archive (Fig. 9.7).

Figure 9.7 Age-depth model for ENC02. The upper part (above 1906) is based on 210Pb dating by Salvetti (2006). The sedimentation rate changes below 1822.

Event deposits F and G have a combined age of 1892. This does not match with any of the earthquakes in the historic earthquake archive. There are two possible explanations. The first possibility is that they result from an earthquake with a local source, comparable to the Maipo Valley earthquake in 1958. Flores et al. (1960) mention earthquakes in 1850, 1870 to 1880, 1883 and 1905. It is highly probable that these deposits were formed as a result of two of the earthquakes between 1880 and 1905, or of another earthquake in this valley that has not been reported. Lomnitz (1960) mentions the absence of major centers of population. Flores et

95 al. (1960) admits that “little is known about the seismic history of the district”. The other option is that these layers were wrongfully identified as earthquake-triggered turbidites.

Event deposit H is dated at an age of 1835. This coincides with an earthquake that hit Concepción in 1835. The rupture distance is estimated from Fig. 4.5 on which is indicated that the rupture ends offshore of Constitución. The distance used for the calculations is 300 km.

The magnitude is estimated between Ms 8 and Ms 8.5 (Table 9.1). This yields an MMI between 5.6 and 6.2. This value is around the proposed threshold and does not exclude the generation of a turbidite.

Event deposit I has a modeled age of 1822. The best match is an earthquake near Valparaíso in 1822. The rupture distance (175 km) was chosen to be the same as for the other Valparaíso earthquakes in 1985 and 1906. Together with an Ms between 8 and 8.5 this yields an MMI between 6.5 and 7.2. This seems to contradict reports in Lomnitz (1970b) that no damage was reported from San José de Maipo. An MMI of 7- 8 was reported in Santiago, and an MMI between 6.5 and 7.2 was calculated for Lago Lo Encañado (Table 9.1). There must have been high intensities and therefore damage in San José as it lies between Santiago and the lake. This might be an example of the claim that Chilean journalists tend to underestimate the effects of an earthquake (Lomnitz, 1970b).

The next event deposit, J, was dated at 1800. Like the 1892 event, this does not match with any of the earthquakes described in chapter 5 (Table 9.1). The same conclusion can be drawn: i.e. either this was an undetected, local earthquake, or it is not an earthquake-triggered turbidite.

Event deposit K is dated as 1751. This coincides with an earthquake in 1751 on the Concepción segment. The distance to the rupture zone was estimated using Fig. 4.5. Both Comte et al. (1986) and Melnick et al. (2009) place the northern end of the rupture zone near the town of Pichilemu. The distance used for the intensity calculations is 250 km, and the magnitude Ms 8.5. This yields an MMI of 6.6 (Table 9.1).

Event deposit L has a modeled age of 1746. To our best knowledge there was no major earthquake in Central Chile at that time. As for the events in 1892 and 1800, there are two options. Either this is a local earthquake, or it is a misinterpreted event. The deeper we go back in time, the less is known on earthquakes in the Maipo Valley.

The lowermost event deposit in ENC02, event deposit M, has an age of 1729. This matches with an earthquake in 1730. This is seen as one of the strongest earthquakes in the history of Central Chile, and its presence in this archive was therefore expected. The distance to the rupture zone was chosen at the same location as the other Valparaíso earthquakes, at 175 km distance. The magnitude between Ms 8.5 and Ms 9 yields seismic intensities between 7.2 and 7.9 (Table 9.1).

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9.4.4. Comments on the pre-1900 earthquakes

There are no more turbidites below event M. The base of ENC02 is dated as 1689 according to the age-depth model. This is in agreement with our calculations. The next major earthquake we would expect is the 1647 earthquake (Table 9.1).

This age-depth model was created on the basis of what we expected to see in the core. This is not the most preferable method in scientific research. The results should therefore be treated with caution. It does however hold some arguments in favour of its credibility. Every earthquake with an MMI above VI is present in the lake archive and this did not require the use of unrealistic sedimentation rates.

All but one earthquake with an estimated seismic intensity above 6 on the Modified Mercalli Intensity scale are present in the sediment archive of Lago Lo Encañado. The one earthquake that is missing is the 1850 Maipo Valley earthquake. It is thought to have had a similar epicenter as the 1958 earthquake and an Ms between 7 and 7.5 (Lomnitz, 1970b). This would yield an intensity between 8 and 8.5, which is the highest intensity for this lake. There is no evidence of a turbidite corresponding to an earthquake of this size in core ENC02. The only evidence given for the magnitude and epicenter estimates in Lomnitz (1970b) are the intensity in Santiago, and large rockslides 14 km to the south of San José de Maipo. It is not unthinkable that the earthquake had different characteristics than proposed in Lomnitz (1970b). The magnitude and epicenter of local earthquakes is difficult to estimate based on intensity reports of just two places. The unidentified turbidites can be either interpreted as local earthquakes, too small in magnitude to make it into the history books but close enough to the lake to produce a significant seismic intensity. Or they can be explained by a non- seismic trigger. The two largest unidentified turbidites (dated at 1892 and 1800) happened after a long period of seismic inactivity. Before 1800, the earthquake of 1751 and 1746 produced only minor turbidites. The first large earthquake before 1800 was in 1730. This means that there was a gap of 70 years with no major earthquakes. The same is true for the 1892 turbidite. There were no turbidites between 1835 and 1892, representing a gap of 57 years. The instability on the delta must have been high at these moments. Maybe this lowered the threshold, a similar effect as the lake-level drop in 2012. Or the delta collapsed under its own weight without a seismic trigger. The absence of the 1850 Maipo Valley earthquake and the large number of unidentified turbidites demonstrates the difficulty of assessing magnitude and epicenter of local earthquakes in uninhabited valleys.

9.5 Comparing Lago Lo Encañado to other lakes in the proximity Other lakes in the Santiago Metropolitan region have been cored for paleoclimatological studies (von Gunten et al. 2009a; von Gunten et al. 2009b; von Gunten et al. 2009c; Vandenberghe, 2012). The position of turbidites in these cores was included in each age model. These turbidites were interpreted as a result of an earthquake, and are therefore seismites.

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In Laguna El Ocho, two seismites were discovered (von Gunten et al. 2009b). The first turbidite was dated as approx. the year 1960, based on a the peak in 137Cs, which corresponds to the Southern Hemisphere fall-out maximum between 1963 and 1965 stratigraphicaly just above the turbidite. The authors link it to the 1960 Valdivia earthquake. The other turbidite was dated at 1850, corresponding with the local earthquake of 1850 in the Maipo Valley. As was the case for Lago Lo Encañado, we believe that the turbidite interpreted as a result of the 1960 earthquake, was in fact caused by the 1958 Maipo Valley earthquake. Using the same formula and epicenter as in the previous chapter, this earthquake produced a seismic activity of 6.8 on the MMI scale (Table 9.2). Both recorded earthquakes were local, intraplate earthquakes. According to this study, no turbidites were triggered by the 1985 or 1906 subduction earthquakes, although the theoretical seismic intensities produced by these earthquakes should have been sufficient (6.2 and 6.8 on the MMI scale respectively).

In Laguna Negra, only one turbidite was described by von Gunten et al. (2009b). It has a dated age of 1906, corresponding to the 1906 Valparaíso earthquake. The core was taken in the south-central part of the lake. Other cores taken closer to the shore and to the delta contain at least six earthquake-related turbidites (Vandenberghe, 2012). These are interpreted as being the result of earthquakes in 1985, 1906, 1850, 1822, 1730 and 1647. In comparison to Lago Lo Encañado, the earthquakes of 1958 and 1835 are missing. On the other hand the earthquake of 1850 was recorded in Laguna Negra, and not in Lago Lo Encañado. There cannot be a distance factor as the lake is located less that 800 m from Lago Lo Encañado.

A core from Laguna Aculeo (von Gunten et al., 2009c) reveals seven shifts in grain-size mode that are interpreted as seismites. The four largest are dated at 1575, 1647, 1730 and 1822. Further seismites are dated at 1906, 1960 and 1985. The seismite interpreted as a result of the 1960 earthquake is again dated by the peak in 137Cs as a result of the Southern Hemisphere fall-out maximum. In Lago Lo Encañado we demonstrated that this turbidite was more likely triggered by the local 1958 Maipo Valley earthquake rather than the 1960 Valdivia earthquake. Laguna Aculeo on the other hand lies closer to the rupture of the 1960 earthquake and further from the epicenter of the 1958 earthquake. The seismic intensity is calculated at 5.3 for the 1960 earthquake and 5.1 for the 1958 earthquake on the Modified Mercalli Intensity scale. According to the theoretical threshold of VI on the MMI scale, neither earthquake should have triggered a turbidite. An alternative explanation is handed by Jenny et al. (2002). These authors interpret the coarse-grained layers as flood deposits rather than earthquake-triggered deposits. A picture of part of a core from Laguna Aculeo is presented in Van Daele (2013). It shows a peak in MS, but no coarse-grained base like we observe in other lakes. Van Daele (2013) does not report any deposits from either the 2010 or the 1960 earthquakes in the lake. This supports the hypothesis in Jenny et al. (2002), but as the coarse- grained deposits seem to coincide with large earthquakes this study cannot exclude the earthquake hypothesis presented in von Gunten (2009c). The presumed seismite dated at 1575 might be also be caused by the 1580 earthquake which caused more damage in Santiago than the 1575 earthquake, but was only discovered in the historical archives in 2012

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(Cisternas et al., 2012). Except for the uncertainty of the 1958/1960 and 1580/1575 earthquakes, all presumed seismites registered in Laguna Aculeo are triggered by subduction earthquakes.

In Laguna del Inca only one seismite was discovered in a core with the oldest date at approx. the year 1900 (von Gunten et al. 2009a). The turbidite is dated at 1906, which corresponds to the 1906 Valparaíso earthquake. This lake is located more than 90 km to the north of Lago Lo Encañado, which explains why apparently only subduction earthquakes of the Valparaíso segment are recorded. Therefore, it is strange that there is no evidence of the 1985 earthquake. We have no knowledge of local earthquakes in that region.

One possible conclusion that can be drawn from table 9.2 is that lakes react to earthquakes in different ways. Some lakes, like Laguna Aculeo (if they are really seismites) and Laguna del Inca, seem to register only large subduction earthquakes. Other lakes, like Laguna El Ocho, only register local earthquakes. Lago Lo Encañado is sensitive to both subduction earthquakes and local earthquakes. It is however difficult to explain this observation. The differences in characteristics (size, depth, catchment) of the lakes discussed above are too strong. For example, Laguna Aculeo and Laguna del Inca show a similar pattern in that they both only recorded subduction earthquakes. However the lakes are completely different. Laguna Aculeo is a very shallow lake (maximum depth of 6 m) with no major slopes surrounding the lake (Van Daele, 2013). Laguna del Inca is a deep lake (maximum depth of > 120 m) at a high altitude (2840 m above sea level) with a mountainous, scarcely vegetated catchment area (von Gunten, 2009a). It is therefore more comparable to Laguna Negra (maximum depth > 350 m, at 2700 m above sea level) (Vandenberghe, 2012). We believe that the fact that only one or another type of earthquake is recorded in a particular lake is only a random effect, most likely as a result of a too small dataset.

A closer look shows that the earthquakes registered in Laguna Aculeo, Laguna del Inca and Laguna El Ocho have a seismic intensity above 6.7, 6.2 and 6.8, respectively. In Laguna Negra the threshold lies above 6.6-6.9 (except for the 1958 Maipo Valley earthquake). We can conclude that the seismic intensity threshold of 6 on the MMI scale, empirically determined in Van Daele (2013) was very accurate in Lago Lo Encañado, but that this is not the case in all lakes. They do, however, fall within the broader threshold (6-7) presented in Monecke et al. (2004). This explanation seems to be more realistic than the previous. It is possible to explain why Lago Lo Encañado has a lower threshold than most other lakes in the region. It is a shallow lake with very steep slopes and a high sediment influx compared to the lake size. The delta in the north and the steep slopes to the east and the west of the lake create an unstable environment in which a turbidite is more easily produced. This threshold can be lowered after a sudden lake-level drop or after a long period of seismic inactivity.

Another explanation for the higher threshold seismic intensities in Laguna Aculeo, del Inca, El Ocho and Negra is the goal of the studies that provided information on turbidites. Vandenberghe (2012) produced a climate reconstruction over the past 500 years for Laguna

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Negra. von Gunten et al. (2009a; 2009b; 2009c) studied anthropogenic pollution and climate changes. These studies all have in common that seismites were not the main interest, and were mainly used to support the age-model. Therefore, it is not unthinkable that smaller turbidites were neglected. Moreover, the lakes were cored at locations at which the influence of turbidites would be the smallest. The intensity thresholds might be lower than presented in this chapter after a paleoseismological study like this study on Lago Lo Encañado.

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Lago Lo Encañado Laguna Negra This study Vandenberghe (2012) year EQ Type Ms/Mw distance (km) MMI distance (km) MMI 2012 Maipo Valley intraplate 4 15 5 15 ## 2010 Maule subduction 8.5 185 7.1 185 7.1 1985 Valparaíso subduction 7.8 175 6.2 175 6.2 1960 Valdivia subduction 8.3 520 4.9 520 4.9 1958 Maipo Valley intraplate 6.3 18 7.3 18 7.3 1943 Illapel subduction 7.9 - 8.3 350 5.1 - 5.7 350 5.1 - 5.7 1939 Chillan intraplate 7.8 - 8.3 340 4.2 - 4.8 340 4.2 - 4.8 1906 Valparaíso subduction 8.3 175 6.9 175 6.9 1861 Mendoza intraplate 7.2 140 5 140 5 1850 Maipo Valley intraplate 7 - 7.5 18 8.1 - 8.6 18 8.1 - 8.6 1835 Concepción subduction 8 - 8.5 300 5.6 - 6.3 300 5.6 - 6.3 1829 Valparaíso subduction 7 175 5.1 175 5.1 1822 Valparaíso subduction 8 - 8.5 175 6.5 - 7.2 175 6.5 - 7.2 1751 Concepción subduction 8.5 250 6.6 250 6.6 1730 Valparaíso subduction 8.5 - 9 175 7.2 - 7.9 175 7.2 - 7.9 1657 Concepción subduction 8 ## ## 380 5.1 1647 Valparaíso subduction 8 - 8.5 ## ## 175 6.5 - 7.2 1580 Santiago ?? ?? ## ## ## ## 1575 La Ligua ?subduction? 7 – 7.5 ## ## ## ##

Table 9.2 Calculated MMI in Lago Lo Encañado, Laguna Negra, Laguna El Ocho, Laguna del Inca and Laguna Aculeo for all historic earthquakes described in chapter 5. Yellow = turbidite present in the core, and MMI > 6. Red = No turbidite, but MMI > 6. Blue = Turbidite present, but MMI < 6. * = Not known which earthquake created the turbidite.

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Laguna El Ocho Laguna Del Inca Laguna Aculeo von Gunten et al. (2009b) von Gunten et al. (2009a) von Gunten et al. (2009c)

year EQ Type Ms/Mw distance (km) MMI distance (km) MMI distance (km) MMI 2011 Maipo Valley intraplate 4 ## ## ## ## ## ## 2010 Maule subduction 8.5 ## ## ## ## ## ## 1985 Valparaíso subduction 7.8 187 6.1 175 6.2 130 6.7 1960 Valdivia subduction 8.3 457 5.2 590 4.7 440 5.3* 1958 Maipo Valley intraplate 6.3 25 6.8 112 5.1 70 5.1* 1943 Illapel subduction 7.9 - 8.3 381 5 - 5.5 270 5.2 - 5.7 342 5.1 - 5.7 1939 Chillan intraplate 7.8 - 8.3 301 4.4 - 5 427 4 - 4.5 293 4.5 - 5 1906 Valparaíso subduction 8.3 182 6.8 160 7.1 130 7.4 1861 Mendoza intraplate 7.2 182 5.3 ## ## 213 4.3 1850 Maipo Valley intraplate 7 - 7.5 25 7.5 - 8.1 ## ## 70 5.9 - 6.4 1835 Concepción subduction 8 - 8.5 ## ## ## ## 230 6 - 6.7 1829 Valparaíso subduction 7 ## ## ## ## 130 5.5 1822 Valparaíso subduction 8 - 8.5 ## ## ## ## 130 7 - 7.7 1751 Concepción subduction 8.5 ## ## ## ## 176 7.2 1730 Valparaíso subduction 8.5 – 9 ## ## ## ## 130 7.7 - 8.4 1657 Concepción subduction 8 ## ## ## ## 300 5.6 1647 Valparaíso subduction 8 - 8.5 ## ## ## ## 130 7 - 7.7 1580 Santiago ?? ?? ## ## ## ## ?? > 6,1* 1575 La Ligua ?subduction? 7 – 7.5 ## ## ## ## 150 4.7 – 6.1*

Continuation of Table 9.2

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9.6 Volcanic activity The region around Lago Lo Encañado is volcanically active, being located in the northern part of the Southern Volcanic Zone (SVZ). The three closest volcanoes (Tupingatito, San José and Maipo) (Fig. 4.7) each had several eruption in the last centuries, yet no tephra layers were encountered in the lake sediments. Evidence for volcanic activity is restricted to mostly colourless volcanic glass in low concentrations. The lack of colour is a sign of high Si-content and thus felsic magma. The main reason that there is very little volcanic material in the lake is that there are no volcanoes in the catchment area. All known eruptions are listed below (Table 9.3). They have a low Volcanic Explosivity Index (VEI) of 2 or lower. Only with very favourable wind conditions (i.e. from the East) ash deposits are expected in this lake. However, the prevailing wind direction in this region is South to South-West (Grass & Cane, 2008). This places Lago Lo Encañado upwind of the three volcanoes. It is therefore not illogical that we observe no evidence of these eruptions.

Table 9.3: List of all known eruptions of the Tupungatito, San José and Maipo Volcano, All in the Maipo Valley.

Tupungatito San José Maipo Year VEI Year VEI Year VEI 1987 2 1960 2 1912 2 1986 1 1959 2 1905 2 1980 2 1895 - 1897 2 1829 2 1968 2 1889 - 1890 2 1826 2 1964 2 1881 2 1961 2 1838 1 1960 2 1822 - 1838 2 1959 2 1959 2 1958 2 1946 2 1925 2 1907 2 1901 2 1897 2 1889 - 1890 2 1861 2 1829 2

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10. Conclusion and outlook

10.1. Conclusion During two field campaigns in 2011 and 2012 a total of 9 cores were taken in Lago Lo Encañado (of which 8 were studied). These cores were examined to reveal evidence of past earthquakes. This evidence was found in fining-upwards layers that were interpreted as turbidites. A microscopic analysis shows that these turbidites consist of terrigenous, clastic sediments and as such they do not provide conclusive evidence for being caused by a seismic trigger (Van Daele, 2013). These sediments originate from the delta in the north of the lake and/or from the steep slopes to the east and the west of the lake. These are instable environments in which a turbidite can be generated without a seismic trigger. There are, however, indications that support the hypothesis of a seismic origin, including the observation of several fining-upwards pulses within one turbidite that can only be explained by turbidites from different parts of the lake triggered at the same time. The occurrence of MTD’s below some of the turbidites additionally increases the chance of these deposits being earthquake-triggered. In total 13 event deposits (events A to M) were found in the most complete cores in the central part of the lake. Eight of these were linked to historical earthquakes found in literature. Event B was linked to the large (MW 8.8) Maule earthquake in 2010. Events C, D and E were attributed to earthquakes in 1985, 1958 and 1906, respectively, based on an age-depth model dated with 210Pb by Salvetti (2006). There was no useful age model available below 1906. Therefore the events below this point were tentatively dated with an age-depth model based on average sedimentary rates. Using this method, events H, I, K and M were attributed to the earthquakes of 1835, 1822, 1751 and 1730, respectively. Events A, F, G, J and L could not be linked to any large subduction earthquake or known local intraplate earthquake. This can in part be explained by the limited knowledge of earthquakes in the scarcely populated Maipo Valley. It is therefore not unthinkable that these turbidites were caused by smaller, local earthquakes that have not been described in literature. For example, event A might be linked to an Mw 4 earthquake with an epicenter at less than 15 km from Lago Lo Encañado.

The seismic intensity on the Modified Mercalli index was calculated for every earthquake using equations in Bakun & Wentworth (1997) for intraplate earthquakes and in Barrientos (1980; in Barrientos, 2007) for subduction earthquakes. Although the magnitudes and rupture zones used in these equations have an increasing uncertainty with increasing age, the seismic intensities obtained from these calculations give a good indication of the strength of an earthquake in the research area. All earthquakes that have left an imprint in the sedimentary archive of this lake had a seismic intensity of > 6 on the MMI scale. This is in agreement with an empirical threshold of 6-7 proposed by Van Daele (2013) and Monecke et al. (2004). The only earthquake that should have been included according to our calculations is a local earthquake in the Maipo Valley in 1850. It was, in fact, the earthquake with the

105 highest local seismic intensity of the past 500 years. The magnitude and location of the earthquake is based on damage reports of only two locations, and it is therefore possible that these data are not accurate. Yet imprints of this earthquake can be found in all other lakes in which it should have had an impact, including in Laguna Negra, less than 800 m to the east of Lago Lo Encañado.

The information on other lakes in the proximity of Lago Lo Encañado comes from paleoclimatologic studies in which the position and date of turbidites was indicated. These lakes are Laguna Negra (Vandenberghe, 2012), Laguna Del Inca (von Gunten, 2009a), Laguna El Ocho (von Gunten, 2009b) and Laguna Aculeo (von Gunten, 2009c). We also calculated the seismic intensity for these lakes, using the same method as we used for Lago Lo Encañado. The main conclusion that could be drawn from this is that the intensity threshold of 6 on the Modified Mercalli Intensity scale, which was accurate for Lago Lo Encañado, does not seem to apply for all lakes. The threshold is 6.6 for Laguna Negra, 6.7 for Laguna Aculeo, and above 6.2 for Laguna del Inca. In summary, the threshold is between 6 and 7 for each lake, as determined by Monecke et al. (2004) and Van Daele (2013). We can conclude from this that Lago Lo Encañado is a very sensitive lake to record earthquakes in the Santiago Metropolitan region and is therefore very well suited for further paleoseismic studies. However, it is important to include in such paleoseismic studies also other lakes in the region, in order to be able to distinguish local earthquakes from large subduction earthquakes. Using lakes on a longer transect could even help to distinguish earthquakes on the different segments.

The observation that these lakes react in a different way to subduction or intraplate earthquakes is interpreted as a coincidence. In Laguna Aculeo and Laguna Del Inca only subduction earthquakes are recorded, in Laguna El Ocho all recorded earthquakes are intraplate earthquakes, and Laguna Negra and Lago Lo Encañado both types are recorded. This hypothesis is rejected as the observations can more easily be explained by the supposed higher intensity thresholds. A larger dataset should eliminate these observations.

10. 2. Outlook As Lago Lo Encañado has proven to be a sensitive archive that holds all earthquakes that reach a seismic intensity of > VI, the lake could be used to reveal earthquakes older than the time of the Spanish arrival (1541 AD). It would require deeper cores and accurate dating in order to construct a reliable chronology. The absence of seismic profiles prevents us, however, to give an estimation on the maximum thickness of the sedimentary infill in this lake.

This study has shown that the type of earthquake has no influence on the appearance of a turbidite. This is only influenced by the local seismic intensity and perhaps also by the season in which the earthquake occurred. On the other hand, this study has also shown that a comparison of dated earthquakes in several lakes can give a better view on the extent of an earthquake. A more accurate comparison would require deposits in these lakes to be studied

106 in more detail. The cores should therefore be looked at in a paleoseismological study rather than as a by-product of a climatological study.

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11. References

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12. Attachments

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Figure 42.1 Core pictures (normal and histogram equalization), core log, magnetic susceptibility and gamma-ray density for ENC01. No spectrophotometric data exist for this core.

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Figure 12.5 Core pictures (normal and histogram equalization), core log, magnetic susceptibility and gamma-ray density for ENC02. No spectrophotometric data exist for this core.

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Figure 12.3 Core pictures (normal and histogram equalization), core log, magnetic susceptibility, spectrophotometric and gamma-ray density for ENC03.

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Figure 12.4 Core pictures (normal and histogram equalization), core log, magnetic susceptibility, spectrophotometric and gamma-ray density for ENC05.

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Figure 12.5 Core pictures (normal and histogram equalization), core log, magnetic susceptibility, spectrophotometric and gamma-ray density for ENC06.

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Figure 12.6 Core pictures (normal and histogram equalization), core log, magnetic susceptibility, spectrophotometric and gamma-ray density for ENC07.

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Figure 12.7 Core pictures (normal and histogram equalization), core log, magnetic susceptibility, spectrophotometric and gamma-ray density for ENC08.

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Figure 12.8 Core pictures (normal and histogram equalization), core log, magnetic susceptibility, spectrophotometric and gamma-ray density for ENC09

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Figure 12.9 Legend of the symbols in the core logs.

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