FACULTEIT WETENSCHAPPEN Opleiding Master in de Geologie

The nature of calcareous rhythmites in the Ordovician of Oslo; bed-by-bed study of the Katian Solvang Formation

Wout Salenbien

Academiejaar 2012–2013

Scriptie voorgelegd tot het behalen van de graad Van Master of Science in de geologie

Promotoren: Dr. T. Vandenbroucke, Prof. Dr. J. Verniers Leescommissie: Prof. Dr. M. De Batist, Prof. Dr. S. Louwye

Nature is wont to hide herself

Ηράκλειτος

Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013

NEDERLANDSTALIGE SAMENVATTING

INLEIDING - DE WERELD GEDURENDE HET ORDOVICIUM Het Ordovicium, gesitueerd tussen 485,4 ± 1.9 Ma en 443,4 ± 1.5 Ma (Gradstein et al., 2012), geldt als een van de interessantere periodes in de aardgeschiedenis. Gedurende het Ordovicium werd de Aarde geconfronteerd met diepgaande veranderingen in biodiversiteit, biocomplexiteit, geochemische balans en ingrijpende veranderingen in het klimaat.

De paleogeografie tijdens het Ordovicium werd gekenmerkt door de overgang van het supercontinent Rodinia naar Pangea, het recentste supercontinent. De dispersie van de continenten, die reeds was ingezet gedurende het Neoproterozoïcum (Li et al., 2008), hield gedurende het Ordovicium aan tot de eerste aanzet voor de vorming van Pangea door de collisie van Avalonië met Baltica omstreeks de Ordovicium-Siluur overgang, 443 Ma (Cocks & Torsvik, 2006; Torsvik & Rehnstrӧm, 2003). Gedurende het Ordovicium verschoof Baltica van een positie nabij de zuidelijke pool richting de paleo-evenaar om tegen het einde van het Ordovicium ongeveer 30° zuiderbreedte te bereiken (Cocks & Torsvik, 2005).

De algemene visie op het klimaat tijdens het Ordovicium veranderde van een monotoon warme periode met intens broeikaseffect en een korte, krachtige glaciatie aan het eind, op aanwijzing van isotopische excursies (Munnecke et al., 2010), naar het beeld van een tijdperk met grote veranderingen in oceaan- biosfeer interacties en mogelijks dus ook in klimaat. Het Ordovicium is ook het tijdperk waarin één van de grootste toenames in biodiversiteit voor mariene leven is geregistreerd (The Great Ordovician Biodiversification Event, GOBE), gevolgd door een grootschalige extinctie aan het eind van het Ordovicium die toegeschreven wordt aan de Hirnantiaan glaciatie. Uit zeespiegelreconstructies (Haq & Schutter, 2008) blijkt dat het Ordovicium gekenmerkt werd door de mogelijks hoogste zeespiegelstanden van het Paleozoïcum. Dit, in combinatie met de extensieve dispersie van de landmassa’s, creëerde een grote toename in gebieden ingenomen door continentaal plat en ondiepe zeeën. Indicaties voor intense magmatische en tektonische activiteit, snelle plaatreconfiguraties en grootschalig vulkanische activiteit bouwen mee aan het beeld van het Ordovicium als dynamische wereld (Harper, 2006).

Mede door de dispersie van de landmassa’s was de paleogeografische dispersie hoog en dit creëerde mogelijkheden voor allopatrische en sympatrische speciatie wat resulteerde in een toename in biodiversiteit. Deze samenloop van omstandigheden wordt verondersteld aan de basis gelegen te hebben van de GOBE, die plaats vond op verschillende tijdstippen op verschillende paleocontinenten (Servais et al., 2010). Dit zorgde ook voor de totstandkoming van de Paleozoïsche Evolutionaire Fauna (Sepkoski, 1981). Die fauna werd gekenmerkt door de ontwikkeling van kust-nabije innovaties zoals suspensie-voeding en de inname van de, voordien grotendeels niet-gebruikte, watermassa’s van de oceaan door phyto- en zoӧplankton. Deze toename in biodiversiteit werd beëindigd door de op-één-na grootste massa-extinctie die 85% van alle mariene soorten van de Aarde wegveegde. In de klassieke visie (Brenchley et al., 1994) werd de extinctie verklaard als een dubbele piek met een eerste

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uitstervingsgolf door de glaciatie en zeespiegeldaling en een tweede golf door het plotse stijgen van de zeespiegel bij het afsmelten die een stagnatie van de oceaancirculatie veroorzaakte (Sheehan, 2001b). Recent onderzoek (Harper et al., 2013) wees er echter op dat een afkoeling onvoldoende is om de extinctie te verklaren en dat de aanleiding gevende factoren gecompliceerder zijn. Daarenboven is het klassieke beeld van het Ordovicium onderhevig aan veel kritiek en aan een herziening toe.

EEN VERNIEUWDE BLIK OP HET LAAT-ORDOVISISCHE KLIMAAT Gedurende de laatste jaren zijn een veelvoud aan technieken, afkomstig uit verschillende disciplines van de aardwetenschappen, toegepast op de studie van het Laat-Ordovisische klimaat en de aanzet tot de Early Paleozoic Icehouse (EPI). Het overgrote deel van die studies wijst op een vroegere start van de afkoeling die aanzet gaf tot de start van de EPI dan eerder werd aangenomen. Dit is gebaseerd op argumenten afkomstig uit geochemische studies (Saltzman & Young, 2005; Trotter et al., 2008; Bergstrӧm et al., 2011), Algemene Circulatie Modellen (GCM’s, Hermann et al., 2004; Nardin et al., 2011), sequentie-stratigrafie (Turner et al., 2011, 2012; Loi et al., 2010) en assemblagestudies van fossiel zoӧplankton (Vandenbroucke et al., 2009, 2010). Andere studies tonen aan dat de glaciaties persisteerden tot in het Siluur (Díaz-Martínez & Grahn, 2007; Page et al., 2007).

GEOCHEMISCHE DATA Geochemische proxies zijn gedurende laatste jaren ontwikkeld tot frequent gebruikte technieken om een gedetailleerd beeld te schetsen van het Ordovicium en andere tijdsperiodes. Deze technieken worden toegepast om stratigrafische correlaties te onderbouwen aangezien veel milieuveranderingen hun invloed hebben op de geochemische samenstelling van de oceanen en de atmosfeer. Aangezien de mengtijden voor de meeste isotopen ogenblikkelijk zijn ten opzichte van een geologische tijdschaal, kunnen sommige veranderingen in de isotopische samenstelling van gesteenten en fossielen, onder bepaalde omstandigheden, informatie verschaffen over verandering in milieu en klimaat doorheen de 18 tijd. Trotter et al. (2008) argumenteert aan de hand van δ Oapatiet waarden dat de temperatuur van oppervlaktewaters tegen het Midden- en Laat-Ordovicium overeenkomstig is met de huidige temperatuur van oppervlaktewaters nabij de evenaar. Dit contrasteert met de eerdere visie dat het Ordovicium gekenmerkt werd door een superbroeikaseffect. Saltzman & Young (2005) argumenteren op 13 basis van δ C data dat de verhoogde sequestratie van organisch koolstof de atmosferische CO2 deed dalen tot beneden de grenswaarden voor de initiatie van ijskapgroei.

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STRATIGRAFISCHE ARGUMENTATIE Loi et al. (2010) en Turner et al. (2011, 2012) rapporteerden zeespiegelcurves, gebaseerd op sequentiestratigrafische studies, die aanwijzing geven tot cycliciteit in zeespiegelveranderingen die enkel verklaard kunnen worden door glacio-eustasie gezien de amplitude en frequentie van de schommelingen. Loi et al. (2010) deden dit voor het Katiaan, maar Turner et al. (2011, 2012) gingen zelfs terug tot het Tremadociaan en Floiaan en argumenteren voor het reeds bestaan van kleine ijskappen in het Vroeg-Ordovicium. Een nadeel van de huidige stratigrafische studies zijn de onzekerheden in datering van de stratigrafische sequenties, wat zijn weerslag heeft op de nauwkeurigheid en correleerbaarheid van de studies, wat dus noopt tot verbetering.

BIOGEOGRAFISCHE RECONSTRUCTIES VAN PLANKTONISCHE ORGANISMEN Gezien de distributie en abundantie van planktonische organismen gecontroleerd wordt door temperatuur, saliniteit en andere factoren, kunnen gegevens omtrent de distributie en abundantie aangewend worden om de verspreiding van planktonische soorten doorheen de tijd te documenteren en zodoende laten ze toe om een reconstructie van de oppervlaktewatertemperatuur te creëren. Dit principe is toegepast door Vandenbroucke et al. (2009, 2010), gebruikmakend van graptolieten en Chitinozoa om een beeld te schetsen van de klimaatsgradiënt van plankton ten tijde van het Sandbiaan. Beide studies besluiten dat de steile gradiënt veel overeenkomsten vertoont met de huidige verdeling van de klimaatgordels en argumenteren voor een koel Ordovicium in het vroeg Laat-Ordovicium. Wanneer deze data verwerkt wordt in Algemene Circulatie Modellen (Herrmann et al., 2004), suggereren deze CO2-waarden van acht maal het huidige atmosferische niveau (8 x PAL; PAL = 280 ppmv). Helaas ontbreken op dit moment gelijksoortige studies voor andere tijdsperiodes gedurende het Ordovicium.

PLAATSING EN MOTIVERING VAN HET ONDERZOEK Elk van bovenstaande methodes suggereert het bestaan van Vroeg- of Midden-Ordovisische ijskappen, maar geen enkele kan op zich gelden als sluitend bewijs. Daarom is er nood aan bijkomend onderzoek, dat zich in deze studie richt op de Oslo-Asker regio in zuidelijk Noorwegen. De Ordovisische successie in Oslo-Asker is grotendeels opgebouwd uit afwisselend mudstone- en kalksteenformaties. Een gedetailleerde beschrijving van de lokale stratigrafie kan teruggevonden worden in Owen et al. (1990). De formatie die van belang is voor dit onderzoek is de ~12 meter dikke Solvang Formatie van midden- Katiaan ouderdom. De Solvang Formatie en andere kalksteenformaties - in feite afwisselingen tussen kalksteenbanken en mudstonelagen - worden door Nielsen (2004), geïnterpreteerd als representatief voor een zeespiegellaagstand in zijn zeespiegelreconstructie voor Baltica (Solvang Lowstand Event). De afwisseling tussen deze kalksteenformaties en tussenliggende mudstoneformaties worden volgens Nielsen (2004) veroorzaakt door glacio-eustasie gezien de amplitude en snelheid van zeespiegelverandering. De kalksteen-mudstone alternaties in de Solvang Formatie worden klassiek gezien als voortgebracht door kortstondige milieu-fluctuaties op een Milankovich- tot

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milleniumtijdschaal en kunnen daardoor dienen als paleoklimatologische indicator. In welke mate ze milieuveranderingen weerspiegelen is echter niet duidelijk en twee uiteenlopende mogelijkheden zijn ofwel een 1-op-1 verandering in depositionele omstandigheden gelijklopend met het klimaat of, volgens Munnecke et al. (2001), zijn deze rhythmieten gevormd door differentiële diagenese van een meer homogeen voorloper-sediment.

DOELSTELLING De verificatie of falsificatie van de interpretatie voor de Solvang Formatie door Nielsen (2004) draagt bij tot het begrijpen hoe kalkrhytmieten gevormd worden en hoe ze geïnterpreteerd kunnen worden in het licht van zeespiegelreconstructies. Daartoe beoogt deze studie het volgende:

- De verificatie van de Solvang Formatie als een laagstand. Dit door biostratigrafische en geochemische karakterisatie en vergelijking met gelijktijdige laagstandafzettingen op andere paleocontinenten. Daarmee zal ook de significantie van de Baltische zeespiegelcurve toenemen (Nielsen, 2004). - De totstandkoming van de kalkrhytmieten van de Solvang Formatie achterhalen aan de hand van een gedetailleerde laag-per-laag monstername en palynologische studie van de Chitinozoa- inhoud van de Solvang Formatie. Parallel zullen de lagen ook geochemisch geanalyseerd worden, in samenwerking met de Friedrich-Alexander Universität Erlangen-Nürnberg, op hun carbonaatgehalte en weinig-mobiele elementen (bv. Ti/Al ratio’s).

Als zich een cyclisch signaal tussen de lagen voordoet, zowel palynologisch als geochemisch, ondersteunt dit de hypothese dat kalkrhythmieten gevormd worden door kortschalige milieufluctuaties. Dit zou ook de interpretatie van de Solvang Formatie als representatief voor een laagstand bevestigen en daarmee credibiliteit verlenen aan de Nielsen (2004) curve. Als er zich anderzijds een uniforme verdeling van de palynomorfen en geochemische data doorheen de lagen voordoet, wijst dit op een diagenetische oorsprong van de alternaties. Tot slot kunnen de resultaten van deze studie vergeleken worden met parallelle studies die, gebruikmakend van eenzelfde methodologie, andere kalkrhytmieten onderzoeken in dezelfde Ordovisische successie (Collart, 2013; Amberg, in prep.).

MATERIAAL Tijdens veldwerk van 1 tot 14 september 2012 werden twee coupes van de Solvang Formatie gedetailleerd opgemeten en bemonsterd, evenals vier andere formaties die in minder detail werden bemonsterd. Dit resulteerde in 179 stalen voor de Solvang Formatie (verdeeld over twee locaties) en een totaal van 293 stalen voor alle bemonsterde formaties, voor later onderzoek aan de Universiteit van Lille 1.

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METHODES & TECHNIEKEN De palynologische studie werd uitgevoerd met behulp van Chitinozoa, een planktonische microfossielengroep met organische wand en onzekere affiniteiten, die aangetroffen wordt vanaf het Midden-Cambrium tot het bovenste Devoon in mariene sedimenten in alle Paleozoïsche oceanen. Doorheen hun tijdspanne kenden Chitinozoa vlugge en goed gedocumenteerde morfologische veranderingen die de groep uiterst waardevol maken voor biostratigrafie in het Vroeg-Paleozoïcum en studies in paleomilieu, paleoklimaat en paleobiogeografie.

De gesteentestalen werden bereid voor palynologische analyse volgens de standaard palynologische methodes, gebruikmakend van HCl en HF. Daarna werden de Chitinozoa handmatig uitgepikt en op glazen plaatjes geplaatst, geschikt voor verdere studie gebruikmakend van elektronenmicroscopie (Scanning Electron Microscopy of SEM). Alle gegevens (determinaties, meetgegevens, staalnummers, etc.) werden gebundeld in een databank waarna verdere statistische en analytische verwerking verwezenlijkt werd met PAST 2.17c (Hammer, Harper & Ryan, 2001).

De geochemische analyses werden uitgevoerd aan de Friedrich-Alexander Universität Erlangen- Nürnberg met het doel de kalksteen- en mudstonelagen geochemisch te karakteriseren. Dit gebeurde aan de hand van de hoofd- en spoorelement ratio’s, die onderverdeeld kunnen worden naargelang hun chemisch karakter. Alle ratio’s werden berekend ten opzichte van aluminium, dat weinig beïnvloed wordt door diagenese. Ti, K, Rb, Si zijn karakteristiek voor kleimineralen, Mg, Fe en Zn voor zowel calciet als kleimineralen en Sr en Mn zijn preferentieel opgenomen in carbonaten. Door de ratio’s van de elementen kan nagegaan worden of diagenese plaatsvond. Deze methode laat echter niet toe om onderscheid te maken tussen differentiële diagenese van een homogeen voorlopersediment en diagenetische versterking van een onderliggend sedimentair rhythmisch signaal.

RESULTATEN

PALYNOGOLISCHE RESULTATEN De analyse van 18 stalen resulteerde in 1636 foto’s waaruit 1612 Chitinozoa herkend werden, waarvan 1535 geïdentificeerd werden tot op genus-niveau en 1209 tot op soort-niveau.

GEOCHEMISCHE RESULTATEN In totaal zijn 19 van de 42 stalen geanalyseerd. Vanuit de resultaten werden de gewichtsprocenten van de individuele elementen en hun ratio’s ten opzichte van aluminium berekend.

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DISCUSSIE

ANALYSE VAN DE PALYNOLOGIE De eerste analyses wijzen op het verschil tussen de vier onderste stalen en de hogere continue sectie. Daarmee wordt besloten dat het gerechtvaardigd is om de onderste vier stalen uit de latere analyses te weren om een stratigrafisch signaal te vermijden. De algemene abundantiematrix zonder de basisstalen toont een uniforme assemblage. Verder uitgewerkte analyses laten toe om 4 soorten te onderscheiden met abundanties die significant typificerend zijn voor een bepaalde lithologie, met name Cyathochitina kuckersiana (kalksteen) en Ancyrochitina onniensis, Belonechitina wesenbergensis brevis en Cyathochitina campanulaeformis (mudstone). Biostratigrafisch wordt de Solvang Formatie gepositioneerd op de overgang tussen de Oandu en Rakvere Stage (Baltoscandisch stratigrafisch system) of de overgang tussen de S. cervicornis en F. spinifera chitinozoa biozone, grotendeels overeenkomstig met de D. clingani graptoliet-biozone.

ANALYSE VAN DE GEOCHEMIE De hoge correleerbaarheid van de elementratio’s impliceert dat er geen grote verschillen optraden gedurende de sedimentatie van de kleien en de samenstelling dus min of meer homogeen was. De geochemische data wijzen op een diagenetische oorsprong van de kalkrhytmieten in plaats van een verandering in het afzettingsmilieu. Dit sluit echter niet uit dat er een onderliggend sedimentologisch signaal verborgen zit in de rhytmieten.

TOTAALZICHT De data kan op twee mogelijke manieren geïnterpreteerd worden. Als men rekening houdt met de bewijskracht van de gehanteerde methoden, kan men besluiten dat zowel de geochemie als de generale palynologische analyses op een diagenetische oorsprong wijzen en slechts een aangepaste abundantiematrix aanwijzing geeft voor een oorsprong in milieufluctuaties. Of men kan meer waarde hechten aan de herkenning van de milieufluctuaties in de weinige soorten die ze vertonen en een mogelijk scenario formuleren dat in staat is alle waarnemingen te verklaren door het combineren van methodes.

Wanneer de data van beide methodes gecombineerd wordt is het mogelijk een ontstaanhypothese te formuleren die door de waarnemingen ondersteund wordt. Chitinozoa-biotopen, hun verspreiding bepaald door oppervlaktetemperaturen, werden reeds gereconstrueerd door Vandenbroucke et al. (2010), waarbij ook een belangrijke grens tussen subtropische en subpolaire biotopen werd geobserveerd rond 35±5° zuiderbreedte. Tijdens de afzetting van de Solvang Formatie lag de Oslo-Asker regio op 30-35° zuiderbreedte, waardoor de chitinozoa-assemblage mogelijks een overgangsbiotoop vertegenwoordigt. Doordat op lagere breedtegraden verschillen in oceanische oppervlaktetemperaturen minder uitgesproken zijn tussen glacialen en interglacialen, is het mogelijk dat deze wisselingen slechts een minimale invloed hebben op de soortenassemblage en het afgezette sediment. Dit zou verklaren waarom slechts een paar soorten een verschil in samenstelling vertonen. De

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eventuele sedimentologische verschillen zouden verborgen zijn door de latere differentiële diagenese. Dit houdt in dat de interpretatie van Nielsen (2004) geldig blijft, hoewel in een minder uitgesproken vorm.

BESLUIT Deze studie had als doel de vorming van de kalkrhytmieten van de Solvang Formatie te verhelderen en zodoende het belang van deze en gelijkaardige formaties te versterken door een beter inzicht in hun vorming te verschaffen. Hoewel de geochemie en algemene palynology beiden argumenteren voor een diagenetische oorsprong, wijst de gefilterde palynologische data op een onderliggend omgevingssignaal. Daardoor kan met enige voorzichtigheid besloten worden dat de kalkrhytmieten voldoen aan de interpretatie zoals deze door Nielsen (2004) is aangewend in de constructie van de zeespiegelcurve van Baltica en dat ze laagstandafzettingen weerspiegelen met hun afwisselende lithologiën, veroorzaakt door milieufluctuaties geïntroduceerd door kortschalige zeespiegelveranderingen, weliswaar diagenetisch versterkt.

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TABLE OF CONTENTS Table of Contents ...... 1 1. Introduction ...... 5 1.1 General setting – The Ordovician World ...... 5 1.1.1 Chronostratigraphy ...... 5 1.1.2 The Ordovician paleogeography – caught in time between two supercontinents ...... 6 1.1.3 Ocean Dynamics in the Ordovician World – Changes in climate, geochemistry and biodiversity ...... 7 1.2 Towards a new understanding of the Late Ordovician climate ...... 10 1.2.1 Geochemical records ...... 10 1.2.1.1 δ18O Values ...... 11

13 1.2.1.2 δ Ccarbonate values ...... 15 1.2.1.3 87Sr/86Sr values ...... 17 1.2.2 Sedimentological evidence for a cooler Ordovician ...... 19 1.2.3 Getting the bigger picture – Ordovician biogeographic reconstructions ...... 23 1.2.3.1 Graptolites ...... 23 1.2.3.2 Chitinozoa ...... 24 1.2.4 Modeling Ordovician climate with General Circulation Models ...... 27 2. Positioning and motivation of the research project ...... 29 2.1 The Oslo-Asker region ...... 29 2.1.1 General geology of the Oslo-Asker region ...... 29 2.1.2 The Ordovician of Oslo-Asker ...... 31 2.1.3 The Baltic sea level curve ...... 35 2.1.4 Rhythmites in lowstand deposits – Solvang Lowstand Event ...... 37 2.1.5 Scientific Question and Hypotheses of this study ...... 38 2.1.6 Objectives of this study: understanding the nature of the calcareous rhythmites of the Solvang Formation ...... 39 3. Materials ...... 41 3.1 In the field ...... 41 3.1.1 Measured sections of the Solvang Formation ...... 43 3.1.2 Bed-by-bed sampling of the Solvang Formation ...... 46

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4. Methods & techniques ...... 47 4.1 Palynology ...... 47 4.1.1 Palynological analysis ...... 47 4.1.1.1 The chitinozoan group ...... 47 4.1.1.2 Sample preparation for palynological analysis ...... 52 4.1.1.3 Chitinozoa picking ...... 52 4.1.1.4 Scanning Electron Microscopy ...... 53 4.1.1.5 Database and statistical Processing ...... 53 4.2 Geochemistry ...... 57 4.2.1 Theoretical background ...... 57 4.2.2 Selected samples ...... 58 5. Results ...... 59 5.1 Palynological results...... 59 5.1.1.1 Description of encountered species ...... 63 5.2 Geochemical results ...... 74 6. Discussion ...... 75 6.1 In-depth evaluation of the palynological results ...... 75 6.1.1 Species level ...... 77 6.1.1.1 Cluster analysis ...... 77 6.1.1.2 Correspondence analysis ...... 80 6.1.1.3 PCO analysis ...... 81 6.1.1.4 ANOSIM ...... 82 6.1.2 Genus level ...... 83 6.1.2.1 Cluster analysis ...... 83 6.1.2.2 Correspondence analysis ...... 84 6.1.2.3 ANOSIM ...... 85 6.1.3 Pushed analyses to tease out maximum environmental signal ...... 85 6.1.4 Exploring variability within one layer ...... 92 6.1.5 Conclusions for the palynological data ...... 93 6.2 Implications for the palynological data ...... 94 6.3 Biostratigraphy ...... 96 6.4 Geochemistry ...... 99

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6.5 An integrated view of the data ...... 102 7. Conclusion ...... 104 8. References ...... 107 9. List of attachments ...... 115

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013

1. INTRODUCTION The Ordovician, ranging from 485.4 ± 1.9 Ma to 443.4 ± 1.5 Ma (Gradstein et al., 2012), arguably stands as one of the more interesting periods in Earth’s history. During the Ordovician, Earth has experienced large systemic shifts in its biodiversity and biocomplexity, global turnovers in its geochemistry and large shifts in the climatic system. A brief review of the major events occurring during the Ordovician is given below from the perspective of plate tectonics, biodiversity and geochemistry and how all these interrelated topics can yield information on the paleoclimate that typified the Ordovician world.

1.1 GENERAL SETTING – THE ORDOVICIAN WORLD

1.1.1 CHRONOSTRATIGRAPHY The earliest recognition of the Ordovician as a system can be attributed to Charles Lapworth, who was the first to identify the need for a separate unit in between the Cambrian and the Silurian systems in 1879. However, the Ordovician only became an official system in 1960 after ratification by the International Geological Congress and much has changed since. The chart in Figure 1.1 shows the current agreement on the Ordovician series and stages.

FIGURE 1.1: THE ORDOVICIAN CHRONOSTRATIGRAPHIC CHART. THE CHART SHOWS PROPOSED CORRELATION BETWEEN THE NEW GLOBAL SERIES AND GLOBAL STAGES AND REGIONAL CHRONOSTRATIGRAPHIC UNITS RECOGNIZED IN THE MAJOR OUTCROP AREAS OF ORDOVICIAN ROCKS. ALSO SHOWN ARE THE STRATIGRAPHIC POSITIONS OF THE STAGE SLICES (SS), DEFINED IN GRADSTEIN ET AL. (2012). FIGURE FROM GRADSTEIN ET AL. (IBID.).

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1.1.2 THE ORDOVICIAN PALEOGEOGRAPHY – CAUGHT IN TIME BETWEEN TWO SUPERCONTINENTS The paleogeography of the Ordovician world is typified by the transition from one supercontinent – Rodinia – to Pangea, the most recent supercontinent known in Earth’s history. The continental break-up of Rodinia, already started in the Neoproterozoic era (Li et al., 2008), still continued in the Ordovician until the first steps towards the formation of Pangea occurred near the Ordovician-Silurian transition 443 Ma ago with the collision of Avalonia and Baltica (Cocks & Torsvik, 2006; Torsvik & Rehnstrӧm, 2003).

Over the last few decades, scientists have come to a general agreement on the large-scale Ordovician paleogeography by compiling evidence from tectonic, paleomagnetic, paleontological and sedimentological sources. The sedimentological evidence for a polar ice cap and its location on Gondwana serves as an important contributor to this agreement, providing ground-truth for the paleomagnetic data and the position of the South Pole. This has resulted in quite similar global reconstructions by several groups of authors, Scotese and McKerrow (1990), Cocks & Torsvik (2006) and Blakey (2008), of the location of the major paleocontinents Laurentia, Siberia, Baltica and the extensive paleocontinent Gondwana. Figure 1.2 depicts their locations and movements throughout the Ordovician (Cocks & Torsvik, 2006). As the area of interest for this study is located in Norway, a particular interest is held in the position of Baltica throughout the Ordovician. Starting in the early Tremadocian (~480 Ma), i.e. the Early Ordovician, it is clear that most of the landmass at that time was located in the southern hemisphere with the South Pole on Gondwana. Baltica was separated from Gondwana and Siberia on its southern to south-eastern border by the Tornquist Ocean and from Laurentia on its west by the Iapetus Ocean, which was at its largest extension at around that time. Avalonia was still attached to Gondwana (present-day South America, Antarctica, Africa, Australia, India and smaller (micro)continents such as North and South China) during the Tremadocian; however rifting had already started and by the end of the Tremadocian Avalonia was separated from Gondwana (Cocks & Torsvik, 2006).

During the Darriwilian (~460 Ma) the development of the Rheic Ocean was the main event, separating Gondwana in the south from Avalonia, Baltica and Laurentia in the north (Cocks & Torsvik, 2006). It is unclear if the rifting and opening of the Rheic Ocean was related to the same tectonic event that resulted in the closure of the Iapetus. The Tornquist Ocean was being subducted beneath Avalonia throughout the Ordovician. By the end of the Ordovician (443 Ma), the Tornquist Ocean was fully closed with Avalonia docking into Baltica while the subduction and closure of the Iapetus Ocean was still ongoing. Throughout the Ordovician, Gondwana moved to the south causing the South Pole to be located more inland at the end of the Ordovician compared to the Early Ordovician, whilst – according to Cocks & Torsvik (2005) - Baltica rotated anti-clockwise from the mid-Cambrian to the end of the Ordovician, meanwhile moving towards the paleoequator (Cocks & Torsvik, 2005), as can be seen on Figure 1.2.

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FIGURE 1.2: THE PLATE-TECTONIC CONFIGURATION OF THE ORDOVICIAN WORLD. FROM LEFT TO RIGHT; TREMADOCIAN (480 MA), DARRIWILIAN (460 MA) AND HIRNANTIAN (440). THE APPROXIMATE POSITION OF THE STUDY AREA HAS BEEN MARKED WITH A RED DOT. THROUGHOUT THE ORDOVICIAN, BALTICA ROTATES ANTI-CLOCKWISE TO ITS CURRENT POSITION. AVALONIA DETACHES FROM GONDWANA AND COLLIDES WITH BALTICA NEAR THE END OF THE ORDOVICIUM, ALONG WITH THE CLOSURE OF THE IAPETUS OCEAN. FIGURE MODIFIED FROM COCKS & TORSVIK (2006, FIG. 5, FIG. 6, FIG, 7).

1.1.3 OCEAN DYNAMICS IN THE ORDOVICIAN WORLD – CHANGES IN CLIMATE, GEOCHEMISTRY AND BIODIVERSITY The view on Ordovician climate has shifted in the last decade from a climatically monotonous period - a greenhouse period with a very short-lived glaciation at the end - to an era with large changes in ocean- biosphere dynamics, and thus potentially in climate, as inferred from the recorded isotopic excursions (Munnecke et al., 2010). Moreover, the Ordovician time period is host to a major diversification in marine life, known as the Great Ordovician Biodiversification Event (GOBE), and the major end- Ordovician extinction event - commonly linked to the Hirnantian glaciation.

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FIGURE 1.3: STRATIGRAPHICAL SCHEME FOR THE ORDOVICIAN. GLOBAL SERIES, GLOBAL STAGES, TIME SLICES (WEBBY ET AL., 2004), δ13C CURVE (BERGSTRӦM ET AL., 2009), 87Sr/86Sr CURVE (SHIELDS ET AL., 2003), SEA LEVEL CURVE (HAQ & SCHUTTER. 2008), REEF DEVELOPMENT (WEBBY, 2002) AND CLIMATE. FIGURE FROM SERVAIS ET AL. (2010, FIG. 1),

From sea level reconstructions it is clear that the Ordovician was characterized by very high sea levels (Fig. 1.3), considered by some authors as the highest sea levels of the Paleozoic (Haq & Schutter, 2008). The large dispersal of the continents during the Ordovician combined with the high sea levels created extensive, epicontinental seas and thereby caused a large increase in the area of continental shelves – and according to Walker et al. (2002) the greatest extent of tropical shelf areas in the Phanerozoic. Indications for intense magmatic and tectonic activity, rapid plate movements and abundant volcanic activity (e.g. the Millbrig and Kinekulle K-bentonites; Bergstrӧm et al., 2004) providing plentiful of nutrients, help to understand the image of the Ordovician as a highly dynamic world (Harper, 2006). Some authors even invoke the presence of mantle plumes as an influence on climatic and geochemical changes – e.g. for explaining the Upper Ordovician shift in the 87Sr/86Sr-curve and the postulated Katian (pre-Hirnantian) warming period, also known as the Boda event (Barnes, 1996; Fortey et al., 2005).

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FIGURE 1.4: THE ‘SEPKOSKI’ DIVERSITY CURVE OF MARINE INVERTEBRATE FAMILIES THROUGH PHANEROZOIC TIME, DOCUMENTING THE CAMBRIAN, PALEOZOIC AND MODERN EVOLUTIONARY FAUNAS. THE GREAT ORDOVICIAN BIODIVERSIFICATION EVENT AND THE ‘BIG FIVE’ MASS EXTINCTIONS OF MARINE INVERTEBRATES. THE GOBE STANDS OUT AS A MASSIVE RISE ON BOTH CURVES. ECOLOGICAL- EVOLUTIONARY UNITS AFTER SHEEHAN (1996): C1-2: CAMBRIAN; P1-4: PALEOZOIC; M1-3: MODERN. GEOLOGICAL PERIODS, FROM LEFT TO RIGHT: C: CAMBRIAN; O: ORDOVICIUM; S: SILURIAN; D: DEVONIAN; C: CARBONIFEROUS; P: PERMIAN; T: TRIASSIC; J: JURASSIC; C: CRETACEOUS; T: TERTIARY. FIGURE FROM SERVAIS ET AL. (2010, FIG. 3).

Biogeographical differentiation was extreme, creating more opportunities for allopatric and sympatric speciation and therefore a rise in biodiversity, as described by Valentine and Moores (1972). This interplay of factors is assumed to have driven the GOBE at different times on different continents (Servais et al., 2010; Munnecke et al., 2010), forming the basis for the Paleozoic Evolutionary Fauna (Sepkoski, 1981), shown in Figure 1.4, and according to Harper (2006) “the greatest and most sustained interval of diversification of life on Earth”. The Paleozoic Evolutionary Fauna was characterized by communities dominated by epifaunal suspension feeders, some carnivores and detritivores. Tiering became more developed, both above and below the sediment-water interface, and a trend towards greater occupation of the water column can be seen (Sepkoski, 1997). Harper (2006) reports a pattern of near-shore innovation and off-shore expansion during the Ordovician. All this resulted in the oceans no

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 longer being vast, sterile water masses, but becoming filled with phyto- and zooplankton - including larvae and graptolites – and the shelves developed into rich communities, occupying more ecospaces, with 30 of the 216 theoretically possible ecospaces occupied by skeletal organisms compared to the 19 occupied ecospaces in the early and middle Cambrian (Bambach et al., 2007).

The clear upward trend of the GOBE (Figure 1.4) came to a halt in the early Late Ordovician with the onset of the second most important extinction of marine life in the Phanerozoic (Sepkoski, 1981). This drop in biodiversity in many fossil groups may reflect a climatic signal of cooling towards the End- Ordovician Hirnantian glaciation. The classical view, as given by Brenchley et al. (1994), depicts the Hirnantian glaciation as a short-lived cold event in an overall greenhouse Ordovician. The Late- Ordovician extinction event would be directly correlated to the glaciation, causing 85% of all marine species to go extinct (Sheehan et al., 2001a). The two recorded pulses in the extinction - in this view - are related to, firstly, the beginning of the glaciation and its induced sea level drop and, secondly, to the subsequent sea level rise and the stagnation of oceanic circulation (Sheehan, 2001b) caused by the melting of the icecap. Recent research, however, has shown that mere ‘cooling’ might not be enough as a cause for the extinction (Hammarlund et al., 2012; Harper et al., 2013) and that the extinction event may have a more complex causal mechanism(s). In addition, the classic view on the Late Ordovician climate has recently been subject to much discussion and revision.

1.2 TOWARDS A NEW UNDERSTANDING OF THE LATE ORDOVICIAN CLIMATE Contrary to the previous held view of a fast cooling towards the short-lived End Ordovician glaciation, data from stable isotopes (Saltzman & Young, 2005; Trotter et al., 2008; Bergstrӧm et al., 2011), General Circulation Modeling (Hermann et al., 2004; Nardin et al., 2011), sequence stratigraphic analyses (Turner et al., 2011) and fossil zooplankton assemblage studies (Vandenbroucke et al., 2009a, 2010) postulate the start of the Early Paleozoic Icehouse (EPI) to take place much earlier in the Ordovician. Saltzman & Young (2005) argue the start of the EPI to be concomitant with the Guttenberg Positive Carbon Isotope Excursion (GICE) in the Katian Stage of the Late Ordovician. Other workers argue that it even started much earlier (Vandenbroucke et al., 2009a), from as early as the Floian (Trotter et al., 2008; Turner et al., 2011). Sedimentological evidence suggests that the icehouse conditions even persisted into the Silurian (Díaz-Martínez & Grahn, 2007; Page et al., 2007).

1.2.1 GEOCHEMICAL RECORDS During the last years, stable isotope studies have been carried out to aid in a more detailed understanding of the Ordovician. Geochemical proxies have become a widely used tool in Earth sciences and have shown great use in stratigraphic correlations because many environmental changes are reflected in the geochemical composition of the ocean and atmosphere. As the ocean mixing time for most isotopes can be considered short to instantaneous on an Ordovician time scale, changes in the geochemical composition of sedimentary rocks and fossils can, under specific conditions, be used as proxies for reconstructing changing environments and climates through time.

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Stable isotope data, however, only provide a veracious record of such changes depending on the preservation state of the source material being examined. Paleozoic rocks have often experienced considerable diagenesis and even metamorphism, altering the original isotopic signal. Therefore, careful consideration is required when choosing which geochemical proxies should be applied, and which source rocks of fossils can be used. Some proxies have been proven more resistant to diagenetic alteration than others and therefore are preferential when influence of diagenesis is suspected but not confidently assessable. The amount of diagenetic alteration is assessable by a range of methods such as trace element analysis, SEM, cathodeluminescence and EBSD (Electron Backscatter Diffraction) C-axis mapping of calcite crystals.

18 1.2.1.1 δ O VALUES The 18O/16O ratio is a well-known stable isotope system and is frequently used as a proxy in paleoclimatology, based on the fractionation of oxygen isotopes, which is mostly depending on temperature and the fluid composition during precipitation of carbonates. 16O is the most common isotope form (99.8%) whereas 18O is much less abundant. The sensitivity to temperature is in the scale of a change of one per mil in the isotopic ratio, which roughly equals a change in temperature of 4°C (Shackleton, 1987). The isotopic ratio, expressed as δ18O, can also serve as a sea-level indicator as during glacial periods, ocean water worldwide is being enriched in the heavier, less easily evaporated 18O whereas the lighter 16O is stored in snow and ice caps. For the Pleistocene, this is correlated to a one per mil increase in δ18O values for a sea level drop of about 100 meters (Shackleton, 1987). As evaporation and precipitation are two main controlling factors on the isotopic ratio, it is also possible to obtain information on the fluid properties affected by these two processes, such as e.g. salinity, which shows a correlation with the δ18O values. This, however, means that changes in salinity affect the δ18O values of the fluid, even if evaporation and precipitation would be absent. Consequently, δ18O values must be interpreted with great care and accurate knowledge of the possible processes that influenced the δ18O values before actual information can be obtained from them. It is evident that this vulnerability to diagenetic alteration dramatically complicates the use of this proxy in ancient deep time.

18 Earlier studies (e.g. Brenchley et al., 2004; Shields et al., 2003) focusing on δ Ocarbonate values obtained by analysis of low-Mg calcite of brachiopods shells - which are believed to be less prone to diagenetic alteration and therefore are explained by those authors to more reliably reflect the original signal - showing a range of -2‰ to -10‰ V-PDB (Vienna Pee Dee belemnite) in the δ18O composition (Shields et 18 al., 2003). These values, assuming a present-day composition (δ Osea water ~ -1‰ V-PDB) for Ordovician sea water, correspond to sea water temperatures of up to 70°C for the Early Ordovician. This unrealistically high temperature - well above the tolerance temperature for the brachiopods that produced the calcite - was explained by some authors by introducing a different sea water composition for the Ordovician (Shields et al., 2003). A suggested mechanism capable of explaining this change in sea water composition is changes in the proportion of high-temperature processes along mid-ocean ridges 18 18 (Wallmann, 2001). Indeed, δ Osea water is an important factor in the equations that translate δ O values into temperatures; Shields et al. (2003) recalculated their Early Ordovician sea water temperatures to 18 range from 27°C to 32°C using a δ Osea water value of <-3 ‰ SMOW. In contrast, Muehlenbachs (1986)

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18 argues that δ Osea water remains more or less constant over time, invoking a buffer effect by sea water/rock interactions and by continental weathering. According to Trotter et al. (2008), a changing 18 δ Osea water over time is inconsistent with altered seafloor basalt compositions and models of the sea 18 water isotopic budget, resulting in essentially constant δ Osea water values as early as the Archean.

18 Moreover, δ Oapatite values obtained from conodont apatite (Trotter et al., 2008) from 20 stratigraphic horizons from 8 sites of similar tropical paleolatitudes allow to discriminate 4 different temperature stages throughout the Ordovician and early Silurian. The first stage coincides with a sustained cooling from greenhouse sea surface temperatures (SST) of ~42°C to ~28°C over ~25 Myr, similar to modern day equatorial temperatures. The second stage, roughly ~20 Myr, shows climatic stability with modern equatorial conditions from the Mid to Late Ordovician. The third stage corresponds to a rapid drop in temperatures during the Hirnantian glaciation and lastly, the fourth stage shows a return to modern day 18 equatorial temperatures by the early Wenlock. This δ Oapatite record disagrees with the earlier proposed 18 supergreenhouse state and also serves as evidence against detrending of the δ Ocarbonate record invoked 18 as an explanation for the anomalously high, reconstructed SST from δ Ocarbonate data for the Early Ordovician. Conodont (biogenic) apatite shows great promises as a robust temperature archive as it is more stable than biogenic marine carbonates (Trotter et al., 2008): apatite is a diagenetically stable composition because the dense crystalline structure inhibits exchange of oxygen in pore waters with the oxygen bound to the phosphates, which constitutes 90% of the oxygen in the conodonts crown tissue. Furthermore, conodonts show an ecological distribution for both nektobenthic (and pelagic?) organisms, allowing the use of them as a potential tool for reconstructing differences between water masses. They serve as an excellent proxy for sea surface temperature reconstructions since it is evident from Trotter et al. (2008) that a strong paleotemperature signal may be derived from conodont apatite (Wheeley et al., 2012). Nevertheless, the reconstructed Early Ordovician temperatures from Trotter et al. (2008) are at the limit of tolerance for shallow epeiric seas and their faunas that characterized the Ordovician. In addition, Sharp et al. (2000) noted that the isotopic composition of biogenic phosphate might be prone to early diagenetic alteration associated with microbial activity.

Wheeley et al. (2012) recently reviewed the use of conodont bio-apatite for oxygen isotope studies by formulating and testing five hypotheses for variability in conodont δ18O. From this study, Wheeley et al. (2012) conclude that conodonts can be heterogeneous with respect to their δ18O composition in a small number of cases, but variability is low within the elements themselves. This heterogeneity however should not be a drawback for paleotemperature reconstructions as it is common to use averaged value of multiple elements. Another observation made by Wheeley et al. (2012) is the tendency for a skew towards lower δ18O values with an increasing conodont Color Alteration Index (CAI) (Epstein et al., 1977). The CAI is a 1 to 5 index for thermal alteration of conodonts, measured by the progressive and irreversible color change caused by the increasing carbonization of trace amounts with higher temperatures. From the dataset, Wheeley et al. (2012) concluded that already from a CAI > 1, there is already increased variability and a shift to lower δ18O values in conodont oxygen.

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FIGURE 1.5: (A) δ18O VARIATION IN OZARKODINA ELEMENTS (ALL CAI 1) FROM ONIBURY, UK, WHERE THE SAMPLE WAS PROCESSES USING DIFFERENT METHODS. THE NUMBER OF TIMES EACH CONODONT WAS ANALYSED IS RECORDED ADJACENT TO EACH PLOT AS A=N. ABBREVIATIONS; NHL: NO HEAVY LIQUIDS; SPT: SODIUM HETEROPOLYTUNGSTATE; B: BROMOFORM. OPEN CIRCLES ARE PROCESSES WITH FORMIC ACID, CLOSE CIRCLES WITH ACETIC ACID. BROMOFORM CAUSES A SHIFT TOWARDS LARGER VALUES AND MORE VARIABILITY, SPT TOWARDS LOWER VALUES. (B) δ18O VARIATION IN PANDERODUS ELEMENTS (ALL CAI 1) FROM THE SAME ONIBURY SAMPLE, UK. FIGURE MODIFIED FROM WHEELEY ET AL. (2012).

Perhaps the most remarkable observation made by Wheeley et al. (2012) is that “the ion microprobe results suggest that some combinations of acids and heavy liquid separation methods employed to extract the conodonts may have a significant effect on δ18O values in the conodont elements investigated.” This is clearly illustrated in Fig. 1.5. According to the authors, variability up to c. 3.2‰ (equivalent to a temperature range of c. 14°C) can occur within a taxon, largely attributed to the use of formic acid processing. From this it is clear that there are still some methodological issues that remain to be solved before δ18O values derived from conodonts can be used as a reliable, routine climate proxy.

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FIGURE 1.6: SYNOPTIC CHART REPRESENTING DIFFERENT ORDOVICIAN GEOCHEMICAL PROXIES AND THEIR SIMILARITIES AND DIFFERENCES. NOTE THE EXCURSION THE LARGE HIRNANATION REGISTERED BY THE CARBON AND OXYGEN ISOTOPES. THE OXYGEN ISOTOPE CURVES FOR BRACHIOPODS AND CONODONTS HAVE DIFFERENT SCALE BARS, BUT DO RESEMBLE IN OVERAL SHAPE. FIGURE FROM MUNNECKE ET AL. (2010).

18 18 Despite their differences in absolute values, both the δ Oapatite and the δ Ocarbonate values show similar trends through time, as can be seen for the Ordovician and Silurian on Fig. 1.6. This confirms the substantial potential of δ18O values obtained from conodonts as a paleotemperature proxy for the Paleozoic, and as Trotter et al. (2008) pointed out, δ18O values derived from conodonts by microprobe analysis suggest an earlier onset of the cooling throughout the Ordovician with lower temperatures than previously believed. However, the conodonts analyzed by Trotter et al. (2008) originated from sites of tropical paleolatitudes and thus are not sufficient as sole sampling sites for global paleo-climate reconstructions (Vandenbroucke et al. 2009, 2010). As indicated by rather uniformly reconstructed tropical sea surface temperatures throughout the Cenozoic glacials and interglacials, data from low- latitude sites may not represent the full range of sea surface temperature fluctuations encountered through glacial and interglacial cycles. So there is a need for more geographically balanced and well- 18 spread sampling for further δ Oapatite analyses (Vandenbroucke et al. 2009, 2010).

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13 1.2.1.2 Δ CCARBONATE VALUES 13 13 12 12 The δ Ccarbonate values refer to the ratio of C to C, with C being the most abundant carbon isotope (98.89%) and 13C much less abundant at only 1.11% of all carbon isotopes. Carbon isotopes are fractionated because of a number of different processes, of which the preferential uptake of 12C during photosynthesis is a substantial factor, therefore depleting the marine organic matter with respect to 13C (c. -25‰ δ13C) according to Hayes et al. (1999). For the dissolved inorganic carbon (DIC) in the ocean surface waters this fractionation results in enrichment in 13C reflecting the phytoplankton production at shallow depth, whereas the deep ocean waters are depleted in 13C because most of the sinking organic matter that has been produced at the surface is remineralized by bacteria (i.e. ‘ocean carbon pump’). 13 According to Cramer & Saltzman (2005), enhanced δ Ccarbonate values in the early Paleozoic can be the result of enhanced burial of organic matter in the deep ocean or by increased productivity at the surface. Unfortunately, Ordovician deep ocean sea floor deposits are not preserved to check this hypothesis (Muller, 2008).

One of the main advantages of the carbon isotopic system compared to the oxygen isotope system is 13 that δ Ccarbonate values are affected to a lesser extent by diagenetic processes. Indeed, most rock systems are closed with respect to carbon since the rocks act as a buffer and pore fluids in rocks tend to contain little carbon (Kaminskas et al., 2010). However, Joachimski (1994) notes that soil-derived fluids from sub-aerially exposed carbonate rocks can alter the isotopic signature. For a similar reason rocks with a carbonatic cement are not suitable for analysis because the unknown origin of the carbon in the 13 cement and the timing of the cementation event(s). The extent of δ Ccarbonate values being reset during meteoric diagenesis can be assessed by a covariance plot with δ18O values (Saltzman & Young, 2005). But it has to be noted that due to the large amount of different processes that can influence the 13 δ Ccarbonate values at the time of their fixation, it is often difficult to correctly interpret them (i.e. more difficult than compared to δ18O data). So care has to be taken in order to deduce meaningful 13 information from δ Ccarbonate analyses.

13 Major changes in the δ Ccarbonate values (isotopic excursions) throughout the Earth’s history might point to major perturbations in the global carbon cycle – changing rates of organic productivity - and can 13 possibly yield information on their causal factors and on the climate or sea level. If δ Ccarbonate values are 13 13 used in conjunction with δ Corg values - the δ C values of organic matter, which is strongly depleted in 13 12 C because the preferential uptake of C during photosynthesis - they can be used as a proxy for pCO2. 13 13 13 Δ C - the paired differential analysis of δ Corg and δ Ccarb - allows for assessing pCO2 since the concentration of dissolved CO2 in the ocean water has a substantial influence on photosynthesis (Kuhn, 2007).

Saltzman and Young (2005) have employed integrated δ13C data in combination with stratigraphical evidence to hypothesize a long-lived glaciation in the Late Ordovician. They describe a positive ~+3% 13 shift in δ Ccarbonate values during the Chatfieldian (Lower Katian) and believe it reflects an enhanced organic carbon burial or oceanic regions with increased production and preservation. The fact that this isotopic excursion is followed by low stand deposits (the Eureka Quartzite) and possible sedimentary gaps might point to the onset of continental ice buildup according to Saltzman and Young (2005). 13 13 Patzkowsky et al. (1997) analyzed both δ Ccarbonate and δ Corganic during the Chatfieldian excursion and

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 reported lowered Δ13C values throughout the excursion, suggesting that the enhanced burial of organic carbon resulted in lowering the pCO2 values to below a CO2 threshold value necessary to allow an ice age to initiate.

FIGURE 1.7: STRATIGRAPHIC DIAGRAM SHOWING CARBON ISOTOPE-BASED CHEMOSTRATIGRAPHIC CORRELATION BETWEEN LOWER KATIAN SUCESSIONS IN THE OSLO REGION AND THOSE IN KENTUCKY, CENTRAL SWEDEN, ESTONIA AND THE YANGTZE PLATFORM IN CHINA. NOTE THE IMPORTANCE OF THE KINNEKILLE AND MILLBRIG BENTONITE FOR REGIONAL CORRELATION. FIGURE FROM BERGSTRӦM ET AL. (2011).

Bergstrӧm et al. (2011) have used δ13C data for the purpose of chemostratigraphy in the Lower Katian, allowing them to assess a global correlation across different paleocontinents and to provide more insight in the sea-level reconstructions for Baltoscandia. The main isotopic event described by Bergstrӧm et al. is the Guttenberg Isotopic Carbon Excursion (GICE), a more widely-used name for the δ13C excursion that Saltzman and Young (2005) called the Chatfieldian excursion. Bergstrӧm et al. (2011) 13 used whole-rock δ Ccarbonate data obtained from the Oslo region and Lake Mjøsa (Norway) and reported two positive δ13C excursions in the Oslo-Asker region, the GICE – located within the Frognerkilen Formation – and the Kope δ13C excursion located within the Solvang Formation. Stratigraphic correlations (Fig. 1.7) based on the geochemical evidence are suggested for Estonia, Sweden, North America and the Yangtze platform in China. This demonstrates the usefulness of δ13C values as a stratigraphical tool.

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87 86 1.2.1.3 SR/ SR VALUES In contrast with the previously described geochemical proxies, the changes in 87Sr/86Sr values are long- term fluctuations through time, mainly because of the 2.4 x 106 years residence time of strontium in oceanic waters (Faure, 1986) and the 103 year mixing rate of ocean waters, causing the 87Sr/86Sr ratio to be homogeneous in the Earth’s oceans at any given time. The strontium isotopic system is characterized by four naturally occurring isotopes (84Sr, 86Sr, 87Sr, 88Sr) of which 87Sr is radiogenically produced by the decay of 87Rb. The strontium isotope system is commonly used to assess global tectonic evolution as one of the main processes that act upon the 87Sr/86Sr ratio is subaerial weathering of continental crust, creating input of 87Sr into the oceans. Another process that is acting on the ratio is the early diagenetic alteration of the seafloor, called halmyrolyse, the latter being an input of non-radiogenic Sr. A last contribution comes by the precipitation and weathering of marine carbonates (Faure, 1986). The changes in the ratio of 87Sr/86Sr values are capable of giving an indication of the tectonic regime. Through these phenomena, it is possible to interpret fundamental tectonic and erosional processes and rates by strontium isotope data.

FIGURE 1.8: GENERALISED TIME CHART, SHOWING CLIMATIC, GEOCHEMICAL AND GEODYNAMIC EVENTS. BODA: BODA EVENT; DAP: DAPINGIAN; CFBP: CONTINENTAL FLOOD BASALT PROVINCE; GICE: GUTTENBERG ISOTOPIC CARBON EXCURSION; HICE: HIRNANTIAN ISOTOPIC CARBON EXCURSION. NOTE; “OUR SCENARIO” IN THE FIGURE REFERS TO THE SCENARIO PROPOSED BY LEFEBVRE ET AL. (2010). FIGURE FROM LEFEBVRE ET AL. (IBID.).

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Ordovician 87Sr/86Sr values in Figure 1.8 clearly display a major and quick shift towards lower values (from ~0.7090 to ~0.7080) occurring during the Sandbian and hence either a larger relative input from young volcanic rocks or a lower input from weathering continental rocks (Faure, 1986). This drop in the 87Sr/86Sr ratios was one of the largest ever recorded for the entire Phanerozoic, equal to the total range of strontium isotope excursion throughout the Cenozoic. A straight-forward explanation for a shift of this scale – both temporal as well as in absolute value - is difficult. Two hypotheses are postulated by Qing et al. (1998). The first is that the rapid decline represents a hiatus or is caused by erroneous assignment of ages. The second hypothesis invokes the combined effects of (1) a reduced continental flux due to the lack of major continental orogenic events and/or transgressive events, resulting in fewer areas with less surface area susceptible for erosion and (2) an increased mantle flux related to a superplume event. The observation that during the Sandbian and Katian a major transgressive event is recognized on sea level reconstructions supports this second hypothesis, because high sea levels halt much of the continental erosion. High sea levels may be caused by high spreading rates, with shallower spreading ridges and shallower deep oceans, due to the higher thermal expansion. Barnes et al. (1995) invoke the presence of a superplume to explain the scale of transgression, the absence of magnetic reversals for much of the Ordovician and the observed shift towards a more pronounced mantle signature in the strontium isotope ratio. The 87Sr/86Sr ratio only began to rise again in the Silurian, presumably because of the higher continental input after deglaciation and the onset of the formation of Pangea.

Lefebvre et al. (2010) have modeled the hypothetical installment of a Large Igneous Province (LIP) caused by a mantle plume and its potential impact on the climate by the concurrent release of atmospheric CO2. They note that the impact of the LIP installment is an accumulation of atmospheric

CO2, reaching a maximum atmospheric carbon dioxide pressure at the end of the LIP emplacement, resulting in a global temperature rise caused by the greenhouse effect of the CO2. These higher global temperatures, however, lead to increased chemical weathering and thus to an increased consumption of atmospheric CO2, ultimately causing a decline in both atmospheric CO2 and global temperature. The basalts deposited with the LIP weather 5 to 10 times faster than granites (Dessert et al., 2003). In the modeling experiment, the atmospheric CO2 level dropped below the initial value - prior to the LIP placement - to a minimum at about 2 Myr after the LIP installment and reached a new, lower, steady state at ~4 Myr after the LIP installment. Lefebvre et al. (2010) further hypothesized that if the installment would have been somewhere in the mid Katian, the expected warming would correspond with the observed, short-lived, end Katian Boda warming event, followed by further cooling into the Hirnantian. Lefebvre et al. (2010) were however careful in their conclusions and pointed out that the model does not allow to confirm that the modeled 0.8 °C warming explains the observed migration of warm-water fauna towards higher latitudes.

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1.2.2 SEDIMENTOLOGICAL EVIDENCE FOR A COOLER ORDOVICIAN One of the most important records available for study of the Earth’s history are the very rocks themselves that have been deposited or that have been altered by the processes that are to be studied. A great deal of information can be inferred from the type of rock that has been deposited, be it in situ or transported from a source region, how it has been deposited and its relation to the other deposits in the stratigraphic succession as well as to contemporary rocks from other localities. Often, sedimentary studies of Ordovician strata are carried out in order to construct sea level curves by compiling information from facies analysis or outcrop-based observations. Among the techniques used, one of the most prominent ones is outcrop-based sequence stratigraphy for basin characterization and the construction of relative sea-level curves. Sequence stratigraphy relies on identifying parasequences and their vertical stacking pattern, revealing patterns of stratigraphical cyclicity. A full sea-level cycle (parasequences) can always be identified by recognizing the strata that coincide with the onset of the base level fall, the end of base level fall, the end of regression and the end of the following transgression (Munnecke et al, 2010). Because sequence stratigraphy is a relatively new way of describing the sedimentary record, there is no consensus on unified terminology available yet, and thus the descriptive names of the horizons may vary according to the researcher that applies them. It goes without saying that this could lead to confusion and hindering easy comparison of different researchers’ work. Therefore a formalization of the terminology by the International Subcommission on Stratigraphic Classification (ISCC) is on its way (Munnecke et al, 2010).

For the correct interpretation of those stacking patterns, one must be aware of the number of different processes than can have similar results but yet have totally different implications when interpreting the parasequences in term of relative sea level variations through time. An example of this is the fact that carbonate and siliciclastic depositional systems response in an opposite way on sea level changes (Munnecke et al., 2010). Furthermore, field-based observations on depositional depth, regime, etc. cannot be used in a straightforward manner to construct sea level curves, but need to be ‘back-stripped’ to compensate for the effect of tectonic subsidence. This is necessary because tectonic subsidence can be excluded as a possible mechanism capable of inducing cyclicity in the sedimentary record (Turner et al., 2011). Further processes that need to be removed are sediment loading and other various smaller effects on the stratigraphy. ‘Back-stripping’ leads to a water-loaded subsidence curve by first removing the effects of compaction and sediment loading from the observed basin subsidence, before correcting for the paleowaterdepth. Cyclicity in sedimentary records can be induced by numerous processes operating at different timescales, ranging from 1st order changes (global tectonics) to 5th order changes on ‘Milankovich’ time scales. Most commonly these processes act on 4th order (autogenic), 3rd order (allogeneic) and 2nd order (regressive-transgressive allogeneic sequence stacking) time scales (Turner et al., 2011).

Loi et al. (2010) studied a sedimentary section in southern Morocco (the Bou Ingarf section), ranging from the early Katian into the Hirnantian, and applied sequence stratigraphy to reconstruct a local sea level curve. They have identified a total of 6 parasequences in the studied section. The only possible explanation for the observed sea level fall events with respect to their temporal range is glacioeustacy. But to verify this mechanism, it must be recognized on a global scale and the observed pattern of sea

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 level rise and fall has to be correlated with the patterns recorded on other paleocontinents. This work is forthcoming for the Hirnantian part of the section and can be compared to other equally complete sections (Ghienne & Vandenbroucke, pers. comm.). In the Katian part of the section, three large abrupt sea level fall events or forced regressions are recognized in the outcrops that can be correlated to Katian events on a global scale. The first of these abrupt sea level falls occurs in the early Katian at c. 453.5 Ma and may correspond with the high-amplitude transgressive-regressive event coinciding with the GICE, both in Laurentia (Eureka Quartzite, Saltzman & Young, 2005) and in Baltoscandia (“Solvang Lowstand Event”, Nielsen, 2004) (Loi et al., 2010). The second abrupt sea level fall takes place in the middle Katian (at c. 450.5 Ma) and can be tentatively correlated with a major regressive event in Baltoscandia identified by Calner et al. (2010) as well as with other events in Laurentia and North Gondwana (Loi et al., 2010). The third abrupt sea level fall occurred at the latest Katian (c. 447 Ma) and can, according to Loi et al. (2010), be related to events reported in the Avalonian domain and possibly in Baltoscandia as the Termi Lowstand Event by Nielsen (2004). Not only sea level falls, but also transgressions and the following highstand periods can be correlated from one locality to another. A good example is the 448- 449 Ma time interval with transgressive, carbonate-prone strata deposited far in the hinterland. This transgression and highstand has temporally been linked with the Boda event, which can be correlated throughout the North Gondwana platform by the growth of carbonate platforms at relatively high latitudes, where they normally are absent (Hammann & Leone, 2007). This contemporaneous sea level rise with increasing temperatures indicates a glacio-eustatic forcing. Yet, high-frequency low-amplitude cycles are maintained throughout this warm event, which means that, according to Loi et al. (2010), the Boda event was not able to melt the intra-continental Gondwana ice cap(s). A major challenge for workers on sequence stratigraphy is now to confirm the suggested correlation of the ‘wiggles’ in the sequence stratigraphic framework (or the deduced sea level curves) using unique and correlatable markers, such as biostratigraphic markers or accurate absolute ages.

Another major study has been carried out by Turner et al. (2011), describing the stratigraphy of the Cape Basin, South Africa. The outcrops are located within the Cape Peninsula and comprise a 500 meters thick Ordovician section, temporally ranging from the Tremadocian to the Darriwilian and containing the Early to Middle Ordovician boundary in between the two members of the Peninsula Formation. The Upper Ordovician was probably removed by the Hirnantian glaciation erosion, as suggested on the Table Mountain by the erosive base of the Hirnantian-aged Pakhuis Formation (Vandenbroucke et al., 2009b), partially overlying Darriwilian aged sediments. The sedimentology of the outcrop comprises thick fluvial sandstones interbedded with thin, subordinate, marine-influenced deposits. They are organized into 10 fining-upward sequences, each 20 to 90 meters thick. The fluvial sandstones have been deposited in extensive braided fluvial systems, and in current days the only locations where such river systems are found, are either in pro-glacial environments or in subtropical semi-arid environments. The major difference between the two environments is the pattern of riverine flow. In the glacial systems they are fed by glacial melt water and thus are influenced by obliquity (40 kyr) and eccentricity (~100 kyr) orbital patterns. According to Turner et al. (2011, p. 1), glacioeustacy provides the only mechanism for the observed sea level fluctuations in excess of 50 meter and in ~1.2 Myr cycles. Moreover Turner et al. (2011) state that the 10 sequences show a thickening and coarsening upward trend in the lower Leeukop Member of the Peninsula Formation and a thinning and fining-up trend in the upper Platteklip

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Member (see Fig. 1.9), being interpreted as “third order allogeneic sequences organized in a second order regressive-transgressive stacking pattern in response to orbitally induced glacial eustacy”.

FIGURE 1.9: THE ORDOVICIAN STRATIGRAPHY WITH THE HAQ & SCHUTTER (2008) SEA LEVEL CURVE AND ONLAP-OFFLAP CURVE AND THEIR PROPOSED CORRELATION WITH THE SEQUENCE STRATIGRAPHY IN THE PENINSULA FORMATION. NOTICE THE STACKING RESULTS IN A GENERAL REGRESSION IN THE LEEUKOP MEMBER AND A TRANSGRESSION FOR THE PLATTEKLIP MEMBER. FIGURE FROM TURNER ET AL. (2011).

The major implication of the mechanism invoked to explain the sea level changes by Turner et al. (2011) is that there already must have been Early Ordovician (Floian) polar ice with a Polar Front at ~40° Southern latitude. Since the scale of the sea level fluctuations is considerably less important compared to the Hirnantian Glaciation, the Early Ordovician glaciations may have been an amalgamation of small ice sheets, which combined, would have a maximum volume of 8 – 12 x 106 km3 to result in a 20 to 30 meter sea level fall. The waxing and waning of those ice sheets in the source area would be the primary control mechanism on the sedimentation in the Peninsula braidplain. Turner et al. (2011) also hypothesize that the absence of glacio-terrestrial deposits of Early Ordovician age is most likely caused by subsequent glacial erosion during later glaciations throughout the Ordovician.

A more recent study (Turner et al., 2012) suggests sedimentological evidence for another Ordovician icehouse period during the Darriwilian. The recorded cyclicity in the Hiswah-Dubaydib Formations in Jordan is considered to have been primarily controlled by eustacy, ranging in amplitudes of tens of meters in a ~1.2 myr cycle and a ~2.4 myr cycle as derived by the Haq & Schutter (2008) sea level curve. By a more modern analogue, Middle Eocene and Holocene ice sequences correspond to ~1.2 long term myr obliquity cycles, while Mesozoic greenhouse cycles show some relation to ~2.4 myr long term eccentricity cycles. If the analogue holds, the changes with amplitude of tens of meters in the

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Darriwilian are problematic in a supposedly ice free, greenhouse world view. However, glacioeustacy is the only known mechanism capable of producing such large and rapid changes. Therefore, Turner et al. (2012) explain the ~1.2 myr cyclic changes observed in the Dubaydib outcrops (member 1 and 2 out of 3) as 3rd order global, glacio-eustatic sequences comprising 4th order, 405 kyr long eccentricity cycles. The ~2.4 myr 3rd order cycles in the Hiswah and Dubaydib (member 3) outcrops are indicative of greenhouse conditions. Turner et al. (2012) further note that their hypothesis collides with the results of other studies, one being that modeled Ordovician atmospheric circulation patterns for the austral winter indicate that atmospheric boundary zones lay ~10° North of their summer positions due to the strengthening of the South Pole high pressure zone (Turner et al., 2012). These northerly shifts have also been observed on seasonal, orbital and geological time scales for the transition to and during the Hirnantian Glacial Maximum (Vandenbroucke et al., 2009a, 2010). Turner et al. (2012) therefore propose the presence of polar ice as early as the Darriwilian (Darriwilian Ice Age), resulting in a Polar Front at ~40° South as indicated by zooplankton provinciality (Vandenbroucke et al., 2010a). An estimate of the maximum ice volume by comparison of the results from Cenozoic ice sheet models is 8 – 12 x 106 km3 or equivalent to 20 – 30 meters glacio-eustatic sea level fall. This ice sheet volume would probably been subject to growth during orbitally induced “cold snaps” causing the expansion and amalgamation of ice sheets. According to Turner et al. (2012) this may also help explain the reduced size of the correlative positive carbon isotopic excursion as the sheets may have only debouched into the sea in a few places. Similar to Turner et al. (2011), the lack of concomitant glacio-terrestrial deposits may be explained by subsequent glacial erosion during later glacial stages as the Hirnantian Glacial Maximum.

A major drawback for the sedimentological data however, is the difficulty to accurately date the sediments beyond stage level and thereby setting limits on how well they can be compared to other eustatic events (Simmons, 2012). This also accounts for the studies performed by Turner et al. (2011, 2012). Nevertheless, the coincidence of climatic change and eustasy is suggested by the concurrence of eustatic events and orbital forcing cycles (Immenhauser, 2005). Furthermore there should be caution when comparing different local eustatic curves to the widely used sea level curve of Haq & Schutter (2008) for the Paleozoic, which is based on different sections in the Great Basin (USA). Indeed differences between models may arise depending on the sequence stratigraphic interpretation strategy used to construct the model. Seemingly, Haq and Schutter (2008) did not fully elaborate on this in their publication but it seems to be vital for allowing a reliable comparison between different curves to be made (Simmons, 2012). When using biostratigraphic calibration for the construction of an eustatic curve, the local biozonation may differ from the global standard biozonation, so care has to be taken when comparing locally constructed curves using biostratigraphic data (Simmons, 2012). As a final remark, Lovell (2010) argued that magmatic underplating of the crust can cause relatively rapid and high frequency sea level changes on a regional scale, Artyushkov et al. (2008) argued that 3rd order, small scale sea level changes can also be explained by variations in local crustal uplift and subsidence. This suggests that the full range of possible mechanisms influencing the reconstruction of sea level curves might still not be known in detail and that further refinement is still possible.

In conclusion, it can be remarked that although sedimentological evidence proves to be an important source of information, it is currently not possible to use it as solid evidence of climate change on its own

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 and therefore there is still a strong need of confirmation and enhancement of the sedimentological patterns by other methods and proxies - not in the least by the accurate dating of events, and the creation of independent correlations across wide areas and paleocontinents.

1.2.3 GETTING THE BIGGER PICTURE – ORDOVICIAN BIOGEOGRAPHIC RECONSTRUCTIONS It is known that physical parameters such as temperature, salinity or various other factors can influence the global distribution and abundance of planktonic organisms. This is especially true for sea surface temperature (SST), controlling the global distribution of foraminifera, calcareous nannoplankton and other planktonic organisms (Kucera, 2007). The distributional patterns of planktonic foraminifera preserved in oceanic deposits have therefore been widely used to assess changes in SST throughout the Quaternary and Neogene and therefore reconstruct changes in oceanic water masses (Dowsett, 2007 for an overview). By documenting the spatial distribution of fossil planktonic organisms through time, a proxy method can be constructed - allowing reconstruction of climate belts distribution and SST in deep time and eventually the testing and verification of numerical climate models (Vandenbroucke et al., 2009a).

1.2.3.1 GRAPTOLITES Researchers (e.g. Cooper et al., 1991; Finney et al., 2007) developed paleoecological models from species distribution and abundance data for graptolites, a zooplankton group that was part of the Ordovician fauna. From the paleoecological models it is clear that graptolite distribution is controlled by depth, latitude and water-masses. Vandenbroucke et al. (2009a) applied the paleobiogeographical distribution of graptolite to provide an integrated reconstruction of Ordovician marine biotopes for the Nemagraptus gracilis time slice, representing about 3 to 4 million years from the base of the Sandbian upwards (for a more detailed definition of the time slice, see Vandenbroucke et al., 2009a). The reconstructed graptolitic marine biotopes have been shown to have features in common with the distribution of present-day planktonic foraminiferal provinces and thus allow mapping climate belt distribution.

The localities that have been used for compiling the data matrix of the species’ presence/absence have all been plotted on the reconstructed paleogeography of the gracilis time slice with an error margin of up to 5°. To elucidate any ecological structure within the compiled presence/absence matrix, two quantitative analyses have been applied (for a more detailed overview of the applied analyses, see Vandenbroucke et al., 2009a) and the resultant presence/absence matrix has been simplified by grouping the localities into bins of 5° paleolatitudes to encompass for the paleogeographical error. This allowed the recognition of four high-level groups in the graptolite data, which are considered to be equivalent to high-level ecological groupings or provinces (Vandenbroucke et al., 2009a). Each major group can still be further subdivided in subgroups or so-called ‘biotopes’. By examining the species that made up each of the groups (and subgroups) it is possible to assign an environmental setting for each group on a shelf-to-basin transect, ranging from shelfal species to deeper-water species and it even

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 allowed for recognizing upwelling zones (Finney & Berry, 1997). This facilitated the assessment of the zooplankton gradient during the gracilis time slice. The gradient recorded from the graptolite distribution data was a steep gradient, similar to a present-day zooplanktonic distribution with a similar latitudinal extent for the tropics and subtropics as to the present-day latitudinal extent. This observation suggests a non-greenhouse state for the climate-ocean system during the gracilis time slice and would imply that a cool world was already established by the early Late Ordovician (Vandenbroucke et al., 2009a). Furthermore, Vandenbroucke et al. (2009a) note that the distribution pattern of graptolite biotopes in the studied gracilis time slice gives support to the climate simulations of the OGCMs (Ocean General Circulation Models) as proposed by Herrmann et al. (2004). These pictured the Ordovician world to have high sea levels and modestly elevated atmospheric pCO2 levels boundary conditions.

Furthermore, the Ordovician atmosphere is suggested to boast higher pCO2 to offset a faint young sun

(Rosing et al., 2010), though the biological proxy data here fits best with the OGCM ran with pCO2 levels at the lower end of the spectrum with values between 8xPAL and 15xPAL (whereas the spectrum ranges up to 22xPAL; Herrmann et al. 2004).

1.2.3.2 CHITINOZOA A similar approach as Vandenbroucke et al. (2009a) has been employed in Vandenbroucke et al. (2010a) but with a different group, namely chitinozoans – presumably the eggs of soft-bodied, unknown marine metazoan animals (Paris & Nõlvak, 1999). The organic walled microfossils found in the fossil record occur in Ordovician to Devonian marine deposits and their significance as a stratigraphic tool has been proven thanks to the high-resolution spatial and temporal record of their occurrences (Vandenbroucke et al., 2010a). According to Paris (1996), the highest abundance of chitinozoans is recorded in high latitude, cold water and locally in outer shelf settings. As can be seen on Figure 1.10 (Vandenbroucke et al., 2010a), the majority of chitinozoans in the fossil record are facies independent - at least within the time slice here studied in detail - which makes most authors believe that their parent organisms had a pelagic mode of distribution and hence their paleolatitudinal spread is largely controlled by sea surface temperature – and thus also by paleocurrents (Vandenbroucke et al., 2010a). This mode of life is also confirmed by their occurrence in anoxic shale, some of which are barren of trace fossils or benthic fauna over enormous distances (Page et al., 2007; Vandenbroucke et al., 2010).

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FIGURE 1.10: COMPARISON OF THE ON-SHELF TO OFF-SHELF DIFFERENTIATION FOR CHITINOZOANS AND GRAPTOLITES DURING THE GRACILIS-TIME SLICE BASED ON PRESENCE-ABSENCE DATA IN VANDENBROUCKE ET AL. (2010A). NOTE THE LARGE (57%) FRACTION OF CHITINOZOANS THAT OCCURRE BOTH ON- AND OFF-SHELF, COMPARED TO THE 17% OF GRAPTOLITE SPECIES. FIGURE DERIVED FROM VANDENBROUCKE ET AL. (2010A).

The methods used in the research of Vandenbroucke et al. (2010a) compare largely the same to the methods as used in Vandenbroucke et al. (2009a) and focuses on the same gracilis time slice or its time equivalent in sections barren of Nemagraptus gracilis fossils. The authors compiled a chitinozoan presence/absence matrix from the literature supplemented by their own data. The wide range of localities have been grouped into 5° latitude bins after analyses (for more details, see Vandenbroucke et al., 2010), from which bins it was possible to distinguish different chitinozoan groups and subgroups, of which the subgroups are considered to be equivalent to the biotopes in Vandenbroucke et al. (2010a) as demonstrated in Figure 1.11.

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FIGURE 1.11: SCHEMATIC DISTRIBUTION OF CHITINOZOAN SPECIES WITH NARROW (BLACK) AND WIDE (GREY) LATITUDINAL RANGES FOR EACH BIOTOPE. THE MAP VIEWS THE NARROWEST RANGES. THE MAIN BOUNDARIES IN TH DISTRIBUTION PATTERNS ARE INDICATED, OCCURRING AT 35°S AND 55°-70°S. VERY FEW SPECIES CROSS THE 35°S BOUNDARY, EXCEPT FOR SUPER-GENERALISTS AND SPECIES OF QUESTIONED TAXONOMY. FIGURE FROM VANDENBROUCKE ET AL. (2010A).

From Figure 1.11, it is clear that chitinozoan biotopes are controlled by sea surface temperatures similarly to modern day foraminifera or similarly to graptolites, with the exception that chitinozoans seem to have been largely restricted to southern Subpolar to Polar waters, poleward of ~35° South during the studied gracilis time slice. Two important boundaries can be recognized at c. 35° South and at 55° - 70° South, representing paleo-oceanographic boundaries that coincide closely with the present-day southerly edge of the subtropics and the austral Polar Front in modern oceans. Much similar to the

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 graptolite study of Vandenbroucke et al. (2009a), the chitinozoan biotopes are characterized by a steep equator-pole faunal gradient that is comparable to the present-day (Vandenbroucke et al. 2010). When modeled, hypothetical plankton provinces retrofitted for the Ordovician reflect the distribution patterns of genuine Sandbian zooplankton with respect to sea surface temperature reconstructions by Herrmann et al. (2004) that use GCMs at 8 times PAL (Present Atmospheric Level, 280 ppmv) pCO2. All this suggests that the early Late Ordovician already experienced a much cooler climate than previously believed, despite the somewhat elevated atmospheric modeled pCO2 and that thus the climate cooling towards the Hirnantian must have started before or during the early Sandbian (Vandenbroucke et al., 2010a).

A final note on the biogeographical information can be that although these biogeographical reconstructions exist and provide crucial information helping in understanding the Ordovician world, there is still a strong need for similar studies focusing on different time slices throughout the Paleozoic, more specifically the Early and Mid Ordovician, as well as for further improvement in the number of sampling sites and recording of the fossil content.

1.2.4 MODELING ORDOVICIAN CLIMATE WITH GENERAL CIRCULATION MODELS Only few General Circulation Models (GCMs) on Ordovician climate have so far been compiled and ran. This is mainly because of the large uncertainties that arise when trying to confine the deep-time boundary conditions for the GCMs. Indeed it is not easy to assess e.g. solar insolation, land albedo and land surface, cloud cover, ocean bathymetry, pCO2 and other greenhouse gasses, atmospheric water vapor, volcanic activity and numerous other factors that each have their significant influence on the climate models. Due to the large uncertainties associated with these boundary conditions, it is questionable to what extent these climate simulations truly represent conditions prevalent during the Ordovician. Therefore the GCMs still remain to be thoroughly tested against other climate-proxydata to assess if these models rendered realistic scenarios before anything can be concluded from them.

In a recent study, Nardin et al. (2011) modeled early Paleozoic climate over a period of 90 million years, ranging from the late Cambrian (Furongian, 490 Ma) over the Middle Ordovician (Dapingian, 470 Ma), Late Ordovician (Katian, 450 Ma), Llandovery (430 Ma) to the Early Devonian (Emsian, 400 Ma). For this, a numerical model, GEOCLIM, has been used to evaluate the impact of paleogeography and rock weathering on atmospheric CO2 levels and climate on a multimillion-year level since Young et al. (2009) and Lefebvre et al. (2010) both suggested that weathering of freshly erupted volcanic material may have triggered the onset of the cooling that led towards the Hirnantian Glacial Maximum; whereas Kump et al. (1999) invoke an increase in orogenic events as the process responsible for the cooling. The modeling procedure followed by Nardin et al. (2011) couples a 3D-atmospheric GCM (the fast ocean-atmosphere model FOAM GCM (Jacob, 1997)) to a coupled model of biogeochemical cycles and climate (COMBINE model (Donnadieu et al., 2006b)) and is based on the concept of long-term global climate stability, requiring a balance between continental rock weathering and solid Earth CO2 degassing. This degassing is assumed to be similar to the present-day, mainly because of the lack of solid evidence for a different 12 rate, which is fixed at 6.8 x 10 mol CO2/yr (Gaillardet et al., 1999) assuming that the present-day weathering rate is equal to the degassing rate.

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For each of the five simulated time slices, a paleogeographical reconstruction with the major topographical barriers has been constructed, with the land type set to rocky dessert since there was no substantial plant life at those times. The increase in solar luminosity over time has been accounted for, starting from 95.4% of the present-day value in the Furongian to 96.7% of the present-day value in the Emsian (Crowley et al., 1991). In the model, Nardin et al. (2011) assumed that each continental grid element had the same proportion of outcropping volcanic and shield rocks for a given time slice in the past, with the proportion calculated as a function of time, based on the recorded 87Sr/86Sr ratio (Veizer et al., 1999). A more detailed description of the boundary conditions can be found in Nardin et al. (2011).

The simulations allow the authors to hypothesize that the northward movement of Laurentia, Avalonia,

Baltica and Siberia caused the atmospheric CO2 drawdown between 500 Ma and 470 Ma as these landmasses entered the ITCZ (Inter Tropical Convergence Zone) with increased continental runoff and silicate weathering, promoting increased drawdown of the CO2 from 21 PAL to 13 PAL by 470 Ma. During the Middle Ordovician, CO2 levels remained as low as 11 to 14 PAL, despite the transgressive trend (Haq and Schutter, 2008) that caused less land surface to be susceptible to weathering. These CO2 levels are believed to have been able to sustain a mean annual continental surface temperature of ~16° C, but this effect was counterbalanced by increasing outcrop area of fresh volcanic rocks as suggested by the 87 86 Sr/ Sr curve (see also Figure 1.6). During the Late Ordovician, CO2 drawdown by weathering of fresh volcanic rocks plays a significant role, keeping atmospheric CO2 at low levels, resulting in mean annual continental surface temperatures of ~13.5° C. During the Llandovery, high runoff and weathering cause a further drop in atmospheric CO2 to 8 – 9 PAL by the collision of Baltica, Avalonia and later on Laurentia. During the Early Devonian, a drastic decrease in weathering rates causes a very high atmospheric pCO2 of 22 PAL by the tectonic drift that causes no landmasses to be present around 10° North and north of that - where the model assumes peak runoff to occur. Furthermore, the amalgamation of Baltica and Laurentia causes a continental effect with decreased runoff rates in the 10° South zone, causing the mean annual continental temperature to rise back to around 16° C (Nardin et al., 2011).

Although these models agree with a prolonged Ordovician cooling, Nardin et al. (2011) concluded with the remark that to allow for more precise modeling, a lot of boundary conditions need to be assessed and constrained in more detail since a large number of uncertainties and assumptions may still lower the reliability of the models. And, of course, there is still the need for ground-truthing the models with climate-proxy data.

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2. POSITIONING AND MOTIVATION OF THE RESEARCH PROJECT Despite the many indications or evidence for the presence of a significant Early-Mid Ordovician ice sheet, none of the provided evidence can be held as conclusive. The oxygen isotope studies by Trotter et al. (2008) and others are subject to uncertainties in SMOW composition and, at the present, lack mid/high latitude data. Studies using δ13C data can still benefit from further refinement allowing for more precise chemostratigraphical correlations. The stratigraphical evidence presented by Turner et al. (2011, 2012) still faces problems with accurate dating of the sequences and time intervals during which the sedimentation took place and furthermore still lacks a method for performing independent global correlations between different sections. Paleobiogeographic studies of plankton present convincing evidence for a Sandbian ice sheet, but no studies for the Early and Mid Ordovician are available. Although using GCMs to model the Ordovician world may provide realistic scenarios, it is still uncertain and very difficult to assess to what extent these models really represent the actual conditions, since the uncertainties in constraining the boundary remain large and there still is a lack of proxy data to ground- truth the models. This warrants further investigations as proposed in this study, which will focus on the Oslo-Asker region.

2.1 THE OSLO-ASKER REGION The fieldwork and studied outcrops for this research are situated in the Oslo-Asker region in southern Norway. Below, an overview of the general geology of the region will be provided as well as how it relates to the overview of the Ordovician world already provided. Furthermore, the sea level curve Nielsen (2004) compiled, based on the outcrops found in the Oslo-Asker region, and its implications will be discussed, specifically for the geological formation of interest on which this research focuses.

2.1.1 GENERAL GEOLOGY OF THE OSLO-ASKER REGION The Oslo-Asker region is located within and largely confined by the Permian Oslo Graben (Dons & Larsen, 1978) and is some 220 kilometers long and 40 to 70 kilometers wide. During the Paleozoic, however, the area consisted of a cratonic basin consisting of an eroded Precambrian basement (Worsley et al., 1983) on which a relatively thick Cambrian to Silurian marine succession was deposited. This succession is considerably thicker and more complete compared to other contemporaneous sequences elsewhere on the Baltic craton (Bruton et al., 1985). The Lower Paleozoic succession in the Oslo-Asker region did not experience a great influence of the advancing nappes and obduction slices of the Scandinavian Caledonides until the mid-Silurian (Bruton & Harper, 1988).

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FIGURE 2. 1: SIMPLIFIED GEOLOGICAL MAP OF THE OSLO-ASKER REGION. MODIFIED FROM NILSEN, O. (1997) AND BASED ON NATERSTAD (1991).

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With the advancement of the nappes, the introduction of a new tectonic regime resulted in the thrusting of the whole succession in varying degrees, up to 150 kilometers southwards for the northernmost parts, to at least parautochthonous for the rest of the succession, with the Cambrian aged Alum Shale serving as a décollement horizon (Harper & Owen, 1983). This resulted in varying degrees of tilting and faulting within the sedimentary succession. By the end of the Permian, a failed ocean rifting event resulted in the development of the Oslo Graben, which was accompanied by magmatism and the installment of igneous rocks as well as volcanic deposits on top of Carboniferous and Lower Permian sedimentary rocks that have been deposited after the Caledonian orogeny. Inside the Oslo Graben, the Lower Paleozoic succession has been preserved from substantial erosion by the graben structure; the low grounds in the city center of Oslo are made up of the Lower Paleozoic shale and limestone.

2.1.2 THE ORDOVICIAN OF OSLO-ASKER The Ordovician rocks of the Oslo-Asker region largely comprise alternating shale and limestone units with rare sandstone (except in the uppermost Ordovician) and some volcanic rocks as bentonite layers in the Middle Ordovician. In the Oslo-Asker region, the rocks are tightly folded and locally overturned. Overall they have a SW-NE strike. Considerable faulting is also present due to the Alum Shale acting as a décollement horizon, with many of the faults being listric splays from the sole thrust (Ramberg & Bockelie, 1981). Throughout the Ordovician section, Permian-aged igneous rocks and dykes can be found. A short schematic overview of the different formations comprising the Ordovician of the Oslo- Asker region will be given below, from bottom to top. All information on the Ordovician sections below is derived from Owen et al. (1990; additional information and references therein). A selection of a few outcrop photographs can be found in Figure 2.2. A stratigraphic table displaying the formations in their relative organization can be found in Figure 2.4.

- Alum Shale Formation; extending from the middle Cambrium to the lowest Ordovician, the

Cambrian-Ordovician boundary lies near the top within this black, organic rich and CaCO3-poor shale, enriched in various trace elements. This shale can be recognized over large parts of Baltoscandia. The thickness ranges from about 20 meters to almost 100 meters at some places. - Bjørkåsholmen Formation; a thin (~1 meter to ~1.2 meters thick) nodular limestone which contains dark concretions near its base and at some localities has shale developed near the uppermost layer of ‘glauconitic’ limestone. - Tøyen Formation; a black and grey shale formation with a thickness of ~20 meters in the Oslo- Asker region. In the well-developed and exposed sections, two members can be recognized, namely the lower Hagastrand Member with rusty weathering striped shale followed by the Galgeberg Member with black shale. In the poorly developed Tøyen Formation localities, only the Galgeberg Member is recognized, which might imply a hiatus at its base. The Tremadocian- Floian boundary lies within this formation. - Huk Formation; this tripartite limestone unit consists of two limestone units with a limestone- shale unit in between them. Thickness for the whole formation varies around ~7 meters in the Oslo-Asker region. The lower Hukodden Member is a fine-grained limestone unit with accumulations of shelly material and is locally dolomitized. The middle Lysaker Member consists

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of limestone and calcareous shale alternations that locally are tectonically deformed and repeated by faulting. The upper Svartodden Member is once more a limestone unit, characterized by an abundant occurrence of orthocone nautiloid near the top. - Elnes Formation; a thick shale formation (thickness is estimated between 60 and 120 meters). The formation has been heavily deformed due to its lower competence in comparison with the adjacent formations. In the Elnes Formation, four different members can be recognized. The Elnes Formation is located in the uppermost Middle Ordovician. - Vollen Formation; this unit comprises nodular limestone horizons alternating with calcareous shale, in which - according to Bjørlykke (1974, p. 14) – two cycles may be distinguished with average thicknesses of 10 centimeters and 40 centimeters. They may represent cycles of respectively 15 kyr and 60 kyr. These short cycles in a 1:4 ratio do not fit any Ordovician orbital periodicities, possibly because of differential compaction of the carbonate and shale beds (Svensen et al., in prep.). The total thickness of the Vollen Formation is not accurately known, but outcrops at Bygdøy are 40 meters thick whereas the outcrop in Vollen is 44 meters thick, without having the base of the formation exposed. - Arnestad Formation; this tectonically deformed formation consists of dark shale with bentonite layers at discrete horizons. The thickness is only estimated to be about 22 meters, but cannot be assessed accurately because of the heavy deformation of the rheologically soft shale unit in between two limestone units. Throughout the horizon, small subordinate limestone nodule horizons are developed, commonly less than 10 centimeters thick. The Arnestad Formation contains the Oslo-Asker equivalent of the Kinnekulle K-Bentonite as well as bentonites that can be correlated with the ones found in Sweden and Estonia and even in North America (Svensen et al., in prep.) albeit the bentonites in North America originated in different areas but are closely spaced in time. Svensen et al. (in prep.) also report ‘Milankovich’ cyclicity in the Arnestad Formation, reporting short and long eccentricity in expected positions relative to the obliquity. They do not find a precessional signal, which may be caused by various unknown parameters. - Frognerkilen Formation; this unit consists of bedded limestone and shale with a rubbly appearance. Towards the top of the 10 to 18 meters thick formation it evolves in more bedded limestone (up to 10 centimeters) alternating with shale layers (5 – 30 centimeters). Bergstrӧm et al. (2011) recorded the GICE in the Frognerkilen Formation, therefore dating the Frognerkilen to an earliest Katian age. - Nakkholmen Formation; a 14 meters thick dark shale formation containing isolated ellipsoidal limestone nodules and pyrite nodules. Towards the top of the formation, these nodules become small, distinct horizons. - Solvang Formation; this formation comprises nodular limestone and bedded limestone alternating with shale. According to Bergstrӧm et al. (2011), the Solvang Formation yields the low amplitude KOPE isotopic excursion (not to be confused with the older GICE), therefore assigning a middle Katian age to the formation. - Venstøp Formation; a c. 10 meter thick dark shale, which in places is graptolitic. The shale have a rusty weathering and contain scattered limestone nodules. The base is taken at the earlier mentioned phosphorite conglomerate horizon. Towards the top of the formation, there is a gradual increase in limestone horizons making it difficult to confine the top of this formation.

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- Grimsøya Formation; this limestone gradually develops from the underlying Venstøp Formation, and thus its base consists of rusty weathering shale embedded in limestone horizons. The upper part of the formation evolves towards more tightly nodular limestone with small embedded shale layers. Near the top of the Grimsøya, the first siltstone layers start to occur, indicating a trend to different depositional conditions. - Skjerholmen Formation; this unit has a different lithology compared to the limestone and shale alternations lower in the succession and is characterized by calcareous shale with limestone, siltstone and even fine sandstone. The siltstone and sandstone become more common towards the top within the formation. The boundary between the Skjerholmen Formation and the underlying Grimsøya Formation is gradual and not easily distinguished. - Skogerholmen Formation; the lithology of the formation is made up of limestone, siltstones and shale with a combined thickness of about c. 35 meter. Two members, the Hovedøya Member and the Spannsloket Member, can be distinguished. - Husbergøya Formation; a thick shale unit with progressively more sandstone layers towards the top, which can be up to 5 meters thick near the top. - Langøyene Formation; a thick sandstone unit, with a lot of the sandstone beds to be interpreted as having a storm surge origin with evidence of channeling. This arenaceous unit forms the top of the Ordovician in the Oslo-Asker region.

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FIGURE 2. 2: OVERVIEW OF SELECTED OUTCROPS OF DIFFERENT FORMATIONS. A: HUK FORMATION, WITH THE 3 MEMBERS INDICATED (WOODMILL LOCATION). B: FROGNERKILEN FORMATION, (NAKKHOLMEN). C: PYRITE LAYER NEAR THE TOP OF THE NAKKHOLMEN FORMATION, INDICATING THE BASE OF THE SOLVANG FORMATION ACCORDING TO SOME AUTHORS (NAKKHOLMEN). D: MIDDLE PART OF SOLVANG FORMATION (NAKKHOLMEN). E: UPPERMOST LIMESTONE LAYER OF SOLVANG FORMATION FOLLOWED BY VENSTØP FORMATION (NAKKHOLMEN). F: TRANSITIONAL BOUNDARY BETWEEN VENSTØP FORMATION AND THE OVERLYING GRIMSØYA FORMATION (NAKKHOLMEN). G: DETAIL OF THE NODULAR LIMESTONE OF THE GRIMSØYA FORMATION (NAKKHOLMEN). H: TOP OF SKOGERHOLMEN FORMATION FOLLOWED BY SKJERHOLMEN FORMATION (NAKKHOLMEN).

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2.1.3 THE BALTIC SEA LEVEL CURVE Based on the Ordovician sections of the Oslo-Asker region, Nielsen (2004) compiled a relative sea level curve for the Ordovician of Baltoscandia based on lithological and paleontological evidence from this region (see Figure 2.3). At a quick glance, the curve displays a series of rapid alternations between high and low sea levels, on a large scale similar to the relative sea level curve by Haq & Schutter (2008). When looked at in detail, there is considerable difference between the two curves, but at least a part of that difference may be accounted for by imprecisely dated sections and local factors being at play since the curves are constructed on different paleocontinents with different epicratonic basins, with the main Ordovician reference areas for the Haq & Schutter (2008) curve being sections in Utah (USA) for the Early to early Middle Ordovician, south central Oklahoma (USA) for the latest Middle to Late Ordovician and the Upper Mississippi Valley in North America for the latest Ordovician. While interpreting the lithological record, in a general way, the limestone deposits were seen as lowstand deposits (often associated with lowstand events) whilst the shale units where seen as deposited during highstands (or drowning events) (Nielsen, 2004). This simplified view, however, does not hold for every interpretation and as such, the followed reasoning by Nielsen (2004) and a detailed overview on the compilation of the published relative sea level curve can be found in Nielsen (2004).

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FIGURE 2. 3: SEA LEVEL CURVE FOR THE ORDOVICIAN OF BALTOSCANDIA PLOTTED WITH THE GENERAL UK SERIES AND BALTOSCANDIAN STAGES. ABBREVIATIONS: H.I.: HIGHSTAND INTERVAL. L.I: LOWSTAND INTERVAL. R.E.: REGRESSIVE EVENT. D.E.: DROWNING EVENT. L.E.: LOWSTAND EVENT. T.: TRANSGRESSION. L.: LOWSTAND. H.: HIGHSTAND. D.: DROWNING. FIGURE MODIFIED FROM NIELSEN (2004).

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According to Miller et al. (2005), the only known mechanism capable of explaining the rapid fluctuations, recorded in the lithological succession as alternating thick shale units with calcareous rhythmites, is glacioeustacy - being the only known mechanism capable of producing rapid, high amplitude fluctuations in sea level. This implies the presence of a sufficiently large polar icecap. The range of sea level changes from highstand to lowstand extremes for the Ordovician of Oslo-Asker is up to 250 meter according to Nielsen (2004).

2.1.4 RHYTHMITES IN LOWSTAND DEPOSITS – SOLVANG LOWSTAND EVENT In the relative sea level curve for Baltoscandia by Nielsen (2004), limestone units interpreted as representing lowstands are in reality quite often limestone-marl alternations or calcareous rhythmites. When looked at in detail, these calcareous rhythmites are classically interpreted as displaying short- term fluctuations on a ‘Milankovich’ range to millennial scale and therefore serve as a paleoclimatological proxy. Less clear however, is in what way these fluctuations actually represent environmental changes. One possible view explaining their occurrence could be a one-to-one change in depositional settings along with the climate (Westphal et al., 2004), but according to Munnecke et al. (2001) these calcareous rhythmites can also be created via differential diagenesis by means of dissolution and precipitation of the carbonate in the original sediments.

Bergstrӧm et al. (2011) compiled a chemostratigraphical framework for the Katian of the Oslo-Asker region. This allowed the correlation of the recorded geochemical events with events on other paleocontinents, namely the part of Laurentia that now comprises the North American Midcontinent. By comparing coeval sea level events on both paleocontinents - correlated by recorded excursions in δ13C such as the GICE or KOPE - Bergstrӧm et al. (2011) noted that Nielsen’s (2004) curve for Baltica shows little agreement with the sea level changes recorded on the Northern American Midcontinent. According to Bergstrӧm et al. (2011) this suggests that a considerable part of the sea level changes in the Oslo-Asker region are rather local than eustatic of nature. The only regional Katian highstand that is contemporaneously recorded on both curves is the linearis Drowning Event (see Fig. 2.3, Fig. 2.4), following the Solvang Lowstand Event and therefore possibly confirming the lowstand nature of the Solvang Formation and its contemporaneous deposits (Fig. 2.4).

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FIGURE 2. 4: LOWER TO MID KATIAN SEA LEVEL CURVE FROM THE OSLO REGIO WITH THE GLOBAL AND BALTOSCANDIAN STAGES AND THE LOCAL STRATIGRAPHY. COMPARISON BETWEEN THE INTERPRETATION MADE BY NIELSEN (2004) AND BERGSTRӦM ET AL. (2011), BASED ON δ13C EVIDENCE. ONLY THE LINEARIS DROWNING EVENT OF VORMSI AGE IS INTERPRETED AS EUSTATIC BY BERGSTRӦM ET AL. (2011). FIGURE FROM BERGSTRӦM ET AL. (2011).

2.1.5 SCIENTIFIC QUESTION AND HYPOTHESES OF THIS STUDY Unraveling the nature of the calcareous rhythmites in the Oslo-Asker region is key to understanding the succession in terms of sea level development and its potential as paleoclimatological archive. Here, we focus on the Solvang Formation, which is entirely consisting of calcareous rhythmites. Parallel studies (Collart, 2013; Amberg, in prep.) focus on the other rhythmic formations in the Ordovician section of the study area.

In the Baltic sea level curve by Nielsen (2004), the Solvang Formation is interpreted as representing an important lowstand event - the Solvang Lowstand Event (Nielsen, 2004, p.89). The gradual transition from the Nakkholmen Formation into the Solvang Formation would indicate that this substantial lowering of the sea level proceeded at a relatively low pace but eventually culminated in a significant drop in sea level (Nielsen, 2004). This lowstand is abruptly ended by a major drowning event (linearis Drowning Event) of which the phosphorite conglomerate, marking the boundary between the Solvang

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Formation and the Venstøp Formation (Owen et al., 1990), serves as proof because phosphorite horizons may indicate a sustained period of non-deposition due to drowning.

In this hypothesis, the limestone-mud alternations that are present within the formation would likely represent high-order environmental (or possibly sea level?) fluctuations during a general glacio-eustatic lowstand. The alternative hypothesis however, based on Munnecke et al. (2001) suggests that such calcareous rhythmites can solely form via differential diagenesis and reflect little to no environmental (let alone climatological) signal.

The main scientific question in this study is how to differentiate between these two hypotheses mentioned above. Do these alternations reflect an environmental signal, or are they solely formed via diagenesis?

2.1.6 OBJECTIVES OF THIS STUDY: UNDERSTANDING THE NATURE OF THE CALCAREOUS RHYTHMITES OF THE SOLVANG FORMATION Verifying or falsifying the interpretation of calcareous rhythmites used by Nielsen (2004) in the construction of his Baltic sea level curve by studying the Solvang Formation in detail will help to improve our understanding of calcareous rhythmites and how they should be interpreted when constructing sea level curves.

The objectives for this study are twofold;

(1.) Testing the interpretation of the Solvang Formation as a low-order (i.e. glacio-eustatic) lowstand event: in this case, these deposits must correlate to other contemporaneous low stand deposits on other paleocontinents. Therefore, we will use our own new biostratigraphical and chemostratigraphical data, alongside similar published information, to evaluate this interpretation. We aim to more accurately date the Solvang Formation by use of biostratigraphical information and to correlate it to other stable cratons, in order to help enhance the significance of the Baltic sea level curve of Nielsen (2004). (2.) Unraveling the nature of the calcareous rhythmites in the Solvang Formation. This involves detailed bed-by-bed sampling and a palynological study of the chitinozoan content of the Solvang Formation, thereby effectively characterizing each bed by its content of this enigmatic organic-walled microfossil group. In parallel, the same samples as well as additional samples will be used for the geochemical characterization of the beds. In cooperation with the Friedrich- Alexander Universität Erlangen-Nürnberg, an analysis of the carbonate content of each individual layer as well as of less-mobile elements (e.g. Ti:Al ratios to assess detrital influx, K:Al ratios to assess the clay types, etc.) will be made.

In a bed-by-bed study, a characterization of every bed is possible which, in return, can be compared to the adjacent beds. If a cyclic signal between the beds would arise - either in the palynological signal or in the geochemical signal - this would support the hypothesis that the rhythmic nature of the Solvang

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Formation is caused by changes in the environment on a - typically - ‘Milankovich time-scale’. Alternatively, a uniform distribution of the palynomorphs and the geochemical signal throughout the formation would suggest a diagenetic origin or a cryptic environmental signal. In the latter case, the alternating lithologies do reflect an environmental signal, though one that does not influence the proxy used (in that particular place). By using a multi-proxy approach, this risk is diminished.

If the Solvang Formation would prove to be a genuine calcareous rhythmite with environmentally induced alternations in lithology, this would support the interpretation of being a lowstand deposit during ice house conditions, with high-frequency sea level changes that are glacio-eustatically driven and thus lend extra support to the interpretation made by Nielsen (2004) for the Baltic sea level curve.

Finally, we aim to put in common our results with those of parallel studies using the same methodology to assess the nature of other calcareous rhythmites in the same Ordovician study area (Collart, 2013; Amberg, in prep.).

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3. MATERIALS In this chapter, a short overview on the actual collecting sites on the field will be given, as well as on the measured sections of the Solvang Formation and how the bed-by-bed sampling was performed.

3.1 IN THE FIELD For the aims of this research, fieldwork was carried out during the first two weeks of September 2012 (1/09/2012 to 14/09/2012) by Wout Salenbien and Tim Collart with assistance and guidance by Dr. Thijs Vandenbroucke (Université de Lille 1, France), Prof. Dr. Arne T. Nielsen (Natural History Museum of Denmark) and Dr. Øyvind Hammer (Natural History Museum, University of Oslo) in the Oslo-Asker region in southern Norway. During the fieldwork, an overview of the local geology was obtained during the first few days after which two different outcrops of the Solvang Formation (Fig. 3.1, Fig. 3.2, Fig. 3.3) were measured and sampled in high detail. A more wide-spaced sampling and measurement was performed on the Venstøp Formation, Grimsøya Formation, Skjerholmen Formation and the top of the Nakkholmen Formation on the island of Nakkholmen, to be used in further studies at the University of Lille 1.

FIGURE 3. 1: SCHEMATIC MAP SHOWING THE LOCATION OF THE STUDIED OUTCROP OF THE SOLVANG FORMATION. MODIFIED FROM GOOGLE EARTH

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FIGURE 3. 2: MIDDLE TO TOP OF THE SOLVANG FORMATION OUTCROP IN BYGDØY.

FIGURE 3. 3: DETAIL OF THE ALTERNATIONS OF THE SOLVANG FORMATION AT THE ISLAND OF NAKKHOLMEN.

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The samples taken were each weighing between ~50 and ~500 grams, depending on the condition of the outcrop; the lithology, protected nature of the site (and allowed sample weight on the obtained authorizations) and the ease of deriving the samples from the outcrop.

3.1.1 MEASURED SECTIONS OF THE SOLVANG FORMATION According to the literature (Owen et al. 1990), the Solvang formation has a thickness of c. 12 meters in Rodeløkken and c. 14.5 meters thickness on the island of Nakkholmen. Slightly different thicknesses were measured during the fieldwork (see below). The boundary between the underlying Nakkholmen Formation and the Solvang Formation is somewhat gradual and is taken at a conspicuous pyrite band which is overlain by bedded limestone with subordinate shale in between them. In Bygdøy and on Nakkholmen, the Solvang Formation is developed as irregularly bedded and nodular limestone up to 20 centimeter thick with intervening calcareous shale layers that can be up to 60 centimeter in thickness. The top of the Solvang Formation is drawn at a phosphorite conglomerate that demonstrates a considerable hiatus between the deposition of the Solvang Formation and the overlying unit. For practical purposes however, both studied sections were measured from the bottom of the lowermost thick limestone layer to the top of the uppermost limestone layer.

The first section of the Solvang Formation that was measured in detail is located on Bygdøy on the locality known as ‘Rodeløkken Shore’ (see Fig. 3.1), named after the Rodeløkken Kafe on top of the cliff (N059°54’54.07” E010°41’28.92”). The total thickness of the Solvang Formation as measured was 11.4 meter at this location and a detailed section can be found in Fig. 3.4. Near the bottom of the section, a small beach inlet was eroded into the Solvang Formation resulting in a part of the section missing with an estimated thickness of ~2 meter.

The second section of the Solvang Formation has been measured on the south-eastern border of the island of Nakkholmen (Fig. 3.1), located in the Oslofjord (N059°53’19.03” E010°41’31.52”). The total thickness of the Solvang Formation at this locality was 10.14 meter. A detailed section is given in Figure 3.5.

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FIGURE 3. 4: SECTION OF THE SOLVANG FORMATION AT BYGDØY. FOR FULL DETAIL SEE DIGITAL VERSION FOR MORE DETAIL

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FIGURE 3. 5: SECTION OF THE SOLVANG FORMATION AT NAKKHOLMEN. SEE DIGITAL VERSION FOR MORE DETAIL

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3.1.2 BED-BY-BED SAMPLING OF THE SOLVANG FORMATION Sampling was performed on every bed for the Solvang Formation at both localities. For the shale beds, sampling thickness was typically around 3 to 5 centimeter for each sample. Since the shale beds tend to be up to 60 centimeter thick, two or more samples were taken in the layers more than 10 to 15 centimeters thick, evenly spread out over the thickness of the layer. By taking multiple samples in the thick layers, a more or less constant sampling frequency was maintained throughout the whole section of the Solvang Formation. The limestone beds were typically around 5 to 15 centimeters thick. While sampling the limestone, we endeavored to cover the whole thickness of the layer in one continuous, orientated sample. Orientation (bottom to top) was indicated on the sample with a marker pen. In total, 87 samples were taken at the Bygdøy locality and 92 samples at the Nakkholmen locality for the Solvang Formation.

A total of 293 samples were obtained from all the studied formations; i.e. the Nakkholmen Formation, Solvang Formation, Venstøp Formation, Grimsøya Formation, Skjerholmen Formation and the stratigraphically lower positioned Huk Formation (Hukkoden and Svartodden Members). For each sample, the stratigraphical height was measured and the lithology was noted, of which a record can be found in attachments 9.1.1 to 9.1.5. Only the Solvang Formation was sampled in high detail at two different localities, the other formations were sampled once every meter. The many samples, not used in this Master of Science study, are stored for future research at the University of Lille 1, France.

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4. METHODS & TECHNIQUES The work carried out for this study can be divided in two large parts. One part being the characterization of the calcareous rhythmites – in this particular case the Solvang Formation – by the detailed study of their chitinozoan content. The second part comprises the geochemical characterization and analysis of the limestone and shale beds. Below, a detailed description will follow on how the research was performed on each of the parts.

4.1 PALYNOLOGY Palynology or the study of fine particles in the paleontological record (derived from the Greek παλύνω meaning ‘sprinkle’ or ‘pollen’ and λέγειν or ‘to speak’) is an ideal method for deriving information about the sediments that contain the studied palynomorphs. Sarjeant (2002) provides a narrowed definition describing palynology as “the study of microscopic objects of macromolecular organic composition (i.e. compounds of carbon, hydrogen, nitrogen and oxygen), not capable of dissolving in hydrochloric or hydrofluoric acids.” By this definition, palynology comprises the study of pollen, spores, orbicules, dinocysts, acritarchs, chitinozoans and scolecodonts, but for this study, only chitinozoans will be used.

4.1.1 PALYNOLOGICAL ANALYSIS The palynological analysis consists of identifying the different chitinozoan species that are present in the studied samples as well as their abundance. When each of the studied samples is characterized by means of its chitinozoan content, this will allow comparison to the other samples - in this case the adjacent layers and layers of the same lithology in the calcareous rhythmites of the Solvang Formation. In this study, the assemblage formed by the different chitinozoan species in one layer is being compared to other layers rather than individual chitinozoan species. If the different lithologies - limestone or shale - represent different depositional environments, this should be reflected in the chitinozoan assemblage composition. By this, it is aimed to investigate if different assemblages can be identified and if these assemblages are concomitant with the changes in lithology and thus sea level (and climate). This should be expected as chitinozoan spatial and temporal distribution is controlled by sea surface temperature (Vandenbroucke et al., 2010a). For each of the studied samples, a database was compiled listing all the identified chitinozoans. This allowed grouping all the samples in an overview matrix indicating all the encountered species and their absolute count for each sample. This matrix could then be utilized to verify if the lithologies are characterized by a distinct assemblage.

4.1.1.1 THE CHITINOZOAN GROUP Chitinozoa are an enigmatic organic-walled microfossil group that appeared in the fossil record during the Early Ordovician (Tremadocian) (Paris et al., 1999), although Shen et al. (2013) just recently reported chitinozoans originating from the middle Cambrian (Cambrian Stage 5 or ~510 Ma). Regardless of the first appearance of the Chitinozoa group, which is clearly not well constrained in time yet, chitinozoans are found in the marine sediments from nearly all Paleozoic oceans, either in chainlike structures or

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 isolated in most sedimentary rocks or low grade metamorphic rocks (Paris et al., 1999) until their extinction in the topmost Devonian (latest Famennian). Their abundance ranges from a small number of specimens to several hundred (exceptionally up to several thousand) specimens per gram of rock. The biological affinities of the microfossil group remain unclear and has been the subject of much discussion since the formal introduction of the group by Eisenack (1931), although the most accepted hypothesis at present is that chitinozoans represent the eggs (reproductive bodies) of invertebrate metazoans called chitinozoophorans (Grahn, 1981).

During the years, Chitinozoa have been allotted a range of different affinities, being assigned to the protists, the fungi, Metazoa and as the reproductive bodies of marine animals. The hypothesis that Chitinozoa belong to reproductive stages of marine metazoans was first proposed by Kozlowski (1963), based on exceptionally well preserved clusters of chitinozoans. Kozlowski (1963) however refrained on giving a fixed affinity and stressed that the chitinozoans displayed only a remote analogy with eggs or egg capsules of present-day animals. Further research (Paris, 1981) revealed that no known fossil groups can be matched to the temporal range of the chitinozoans and the main conclusion was that chitinozoans represent the eggs of soft-bodied marine metazoans a few millimeters in length, eventually with a pelagic or nectic mode of life. This may need to be revised in the light of the recent discovery of much older chitinozoans (Chen et al., 2013) that are believed to display a benthic mode of life. Despite the shared belief that chitinozoans represent eggs or egg capsules of marine metazoans, they still occur in three different ways in sediments; as free vesicles, as catenary structures or ‘chains’ (less frequently) and as ‘cocoons’. The link between the different kinds of occurrences has not been fully explained yet (Paris et al., 1999). Due to the widespread occurrence of chitinozoans, a pelagic mode of dissemination of the vesicles is inferred, but the chitinozoophorans would not necessarily have all been pelagic organisms, some of them could have been nectic organisms with the wide paleogeographical distribution resulting from the drifting of the eggs (Paris et al., 1999). Given that the majority of the species are facies independent and that they occur in widespread anoxic deposits that are devoid of trace fossils or benthic fauna and that they are frequently found together with exclusively epipelagic/nektonic organisms, this lends further credibility to the inferred planktonic mode of life (Vandenbroucke et al., 2010a).

Although not observed in this study, all chitinozoans probably had linked vesicles at some stage in their development (Paris & Verniers, 2005) which has been used as an argument for the egg-nature of chitinozoans. Linkage ranges from linear or chain-like structures to planar aggregates. Rarely, chaotic aggregates consisting of singular species are retrieved, and they probably represent faecal pellets.

By accepting Chitinozoa as eggs of marine metazoans, it is inferred they belong to an ontogenetic cycle of organisms and therefore assumed that their morphology is genetically controlled. This allows establishing a suprageneric classification, including orders, families and subfamilies based on morphological criteria. This implies that the observed specific diversity of the chitinozoans reflects the paleobiodiversity of the unknown parent animals, which is something that can be observed in the modern-day fauna and therefore can be extended into the Paleozoic (Paris et al., 1999). Furthermore, analogous to the excellent knowledge of conodonts and scolecodonts and the only fragmented

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 knowledge of the polychaete animals to which they belonged, the systematic affinities are not necessary for the evaluation of the chitinozoan biodiversification pattern and tempo (Paris et al., 1999).

Chitinozoan vesicles comprise a wide range of shapes formed by an organic wall, called a tegument or test (Paris & Verniers, 2005). These shapes are commonly represented in numerous unrelated fossil and extant unicellular organisms or reproductive cycles of metazoans, thus any attempt on assigning biological affinities by the shape of chitinozoans has failed due to the radically different views that were represented (Paris et al., 1999). It is worthy to note that although the name Chitinozoa suggests the composition of the vesicles to be chitinous, no one has been able to prove the presence of chitin in the organic vesicle wall so far (Jacob et al., 2007). This can be explained by (a.) the assumption that the molecular structure of chitin is not preserved through time or (b.) the vesicles actually never contained chitin (Jacob et al., 2007). The vesicle length usually ranges from around 100 micrometer to a few hundred micrometers. In the most basic view, a chitinozoan consists of a chamber closed by a plug (prosome) or an operculum. The chamber may be (hemi)spherical, lenticular, ovoid, conical, cylindrical or claviform (Fig. 4.1).

FIGURE 4. 1: RANGE OF BASIC CHAMBER SHAPE OF CHITINOZOANS. FROM PARIS ET AL. (1999).

On top of the chamber, a circular opening or aperture can be found, either directly on the chamber or at the end of a tube or neck (Fig. 4.2). If the plug is developed within the neck and forms a cylindrical segment with horizontal septa, it is called a prosome, whereas it is called an operculum if it directly seals off the chamber. Anti-aperturally, as an extension of the operculum or prosome, a rica may be developed in various degrees as a membranous extension (Paris & Verniers, 2005). The distinction between operculum-bearing vesicles and prosome-bearing vesicles is a primarily generic classification feature. Another key feature to the classification of Chitinozoa is the shape and ornamentation of the margin, located at the junction between the flanks and the bottom of the chamber (Fig. 4.2). Possible ornamentations of the margin can be the formation of a carina (a circular expansion of the wall around

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 the chamber), processes (appendices or spines with various degrees of complexity and mostly hollow) and intermediate forms (Paris et al., 1999). The vesicles show a wide degree of ornamentation (Fig. 4.3) of the vesicle wall by variation of the size, structure, distribution and shape of these wall ornamentations and this also serves as an important classification factor. Distribution can range from discrete areas being ornamented to geometrical patterns (e.g. rows, rims and ridges) on parts or the whole of the vesicle (Paris & Verniers, 2005). On the apex of the chamber, an anti-apertural mark can frequently be found, ranging from discrete forms as concentric rims, a central scar or central pit, to more developed features as a mucron (membranous tube), copula (long hollow tube), siphon (bulb-like membranous expansion) or a peduncle (solid structure) (Paris & Verniers, 2005).

FIGURE 4. 2: CHITINOZOAN VESICLE WITH MORPHOLOGICAL TERMINOLOGY. A: LAGENOCHITINIDAE. B: DESMOCHITINIDAE. OPERCULUM AND PROSOME DEPICTED IN ORANGE AND YELLOW. FIGURE FROM PARIS ET AL. (2005).

The highest in the suprageneric classification of Chitinozoa is based on the occurrence of an operculum in the Operculatifera and the occurrence of a prosome in the order of the Prosomatifera. The order of the Operculatifera encompasses the Desmochitinidae family, which in turn includes six subfamilies. The order of the Prosomatifera is divided into two families, i.e. the Lagenochitinidae (well distinguishable neck) and the Conochitinidae (absent shoulder, deep prosome). The Lagenochitinidae comprise six subfamilies whereas the Conochitinidae include seven subfamilies. Differentiation on a subfamily-level is largely based on the chamber surface. The classification on a generic level is based on the arrangement, location and development of morphological features as well as on the outline of the vesicle. In total, 56 genera are withheld and presently used. For a classification on species level, the criteria are variations on the morphological parameters that are considered as generic criteria (Paris & Verniers, 2005).

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FIGURE 4. 3: DIFFERENT KINDS OF WALL ORNAMENTATIONS. 1, SMOOTH VERMICULATE; 2, FOVEOLATE; 3, FELTLIKE; 4, SPONGY; 5, VERRUCATE; 6, SIMPLE SPINES; 7, SIMPLE AND BRANCHED HAIRS OF SPINES; 8, BI AND MULTIROOTED SPINES; 9, MESHLIKE STRUCTURE; 10, CRESTS WITH VERTICAL ROWS OF FREE OR CONNECTED SPINES; 11, CRESTS OF WEBLIKE TO DISCONTINUOUS MEMBRANES; 12, COMPLETE OR PERFORATED/RETICULATED CARINA. FIGURE FROM PARIS ET AL. (1999).

Chitinozoans are characterized by rapid morphological changes through time. Combined with the fact that chitinozoans can be retrieved from a large variety of marine sedimentary rocks, their characteristic wide paleogeographical distribution and the fact that they can be studied with a fairly simple technical preparation (more on that below), makes Chitinozoa one of the most useful biostratigraphical microfossil groups for studying the Early Paleozoic Earth. Furthermore, they can provide useful information on paleoenvironments, paleoclimate, paleobiogeography and even post-depositional history of the bearing rocks (e.g. reworking). During the Ordovician, 3 major biozonational schemes for Chitinozoa have been defined; the high-latitude Gondwana assemblages, the moderate-latitude Baltica assemblages and the low-latitude assemblages from Laurentia. This ‘provinciality’ de facto represents SST-sensitivity, with the cited paleocontinents roughly occupying discrete positions in latitudinal climate belts (Vandenbroucke et al., 2010a). Environmental control on Chitinozoa is assessed by various parameters such as abundances (specimens per gram of rock), chitinozoan diversity (species and genera per sample), the relative frequency of the taxa and the ratio of chitinozoans to other elements such as acritarchs (Paris & Verniers, 2005). The shape and preservation of the chitinozoan vesicles can also provide information on the timing of lithification of the sediments they are contained in. Because the organic wall of the vesicles stayed flexible for a considerable amount of time after burial, it is possible to

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 distinguish early lithification if the vesicles are preserved in full relief (e.g. limestone, cherts etc.). In argillaceous sediments the vesicles are often flattened by compaction of the sediments, especially when interstitial water was removed from the sediments (Paris & Verniers, 2005).

4.1.1.2 SAMPLE PREPARATION FOR PALYNOLOGICAL ANALYSIS Sample preparation started by thoroughly cleaning the selected samples by chipping off all weathering surfaces so the sample only consisted of fresh, unweathered rock of a single lithology. After this, all pieces were meticulously scrubbed to remove any dust and any possible source of contamination before being dried at a temperature of 60 °C in an oven for at least 6 hours. The dried samples were physically broken down into small pieces by hammering, after which the fragments were collected in a glass container, approximately 30 grams for limestone samples and ~20 grams for shale samples. Hydrochloric acid (HCl) at 2N strength was added to assess the amount of carbonate in the sample. All processed samples reacted to the HCl, albeit in varying degrees from slightly positive to positive for the shale samples to a strong positive reaction for the limestone samples. This was followed by adding HCl 2N and letting it completely react at 65 °C. Subsequently, the fluid was decanted and replenished with fresh HCl 2N until no reaction occurred anymore. After the HCl was completely reacted - and thus the carbonates removed - the samples were rinsed with pure H2O. Next step involved adding 150 milliliter of hydrofluoric acid (HF) at 40% strength to the sample and placed in a reaction bath at a temperature of 65 °C for 3 consecutive days to remove all the silicates in the sample after which the samples were rinsed with pure H2O by using the decanting method. If there were still small pieces remaining, an additional 100 ml HF 40% was added for one day, kept at 65 °C and followed by one or more cycles of HCl 2N until the whole sample was dissolved after which the sample was rinsed with distilled water by using decantation. Finally, the sample was filtered on a sieve with a mesh size of 53 micrometers. The fraction larger than 53 micrometer was collected for further study whilst the fraction smaller than 53 micrometers was stored separately.

Most of the samples were processed at Ghent University in the laboratory of the Research Unit Paleontology by Sabine Van Cauwenberghe, a detailed list of the weights and details on the processing of the samples prepared at Ghent University can be found in attachment 9.2 . 12 additional samples were processed at the Université Lille 1, France, at the department of Géosystèmes, of which WSA-12- 101B and WSA-12-102 were used in this study.

4.1.1.3 CHITINOZOA PICKING The Chitinozoa were picked out of a watch glass which contained part of the dissolved sample which had been watered down to disperse the material evenly in the watch glass, easing recognition and picking of chitinozoan vesicles. Picking was carried out with the use of handcrafted glass pipettes made out of Pyrex glass, after which the picked Chitinozoa were positioned and concentrated on small glass plates that can be mounted on metal stubs apposite for the SEM. The picking was performed by using an Olympus SZ40 microscope using a magnification of 35 - 40 times.

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4.1.1.4 SCANNING ELECTRON MICROSCOPY After being handpicked, collected on glass plates and mounted on metal stubs, the samples were prepared to be viewed by use of Scanning Electron Microscopy (SEM). As SEM is performed in high vacuum as gas molecules interfere with the electron beam. The Chitinozoa must be prepared for SEM imaging to remain fixed on the glass plate and allow electric conductivity. This was facilitated by using a JEOL JFC1200 sputter coater that disperses a 10 to 20 nanometer thick layer of gold, evenly coating the whole glass plate and thereby reflecting the original surface morphology. This coating with gold has the benefit that the surface is now acting as a conductor rather than an insulator as does the glass and the organic compound of which chitinozoans are made of. This makes it possible to use the electron beam from the SEM for imaging the chitinozoans, as without the conducting behavior of the gold coating, the electrons from the beam would have nowhere to go and cause charging and imaging problems.

The SEM imaging was carried out by using a JEOL JSM-6400 Scanning Microscope along with the Vantage Vista Imaging Software suite v2.3.2 with a beam energy of 12 kV. Chitinozoa were measured on the spot with their longitudinal axis horizontally orientated to prevent any possible error by distortion. After measuring, a picture of every encountered Chitinozoa was made using the Vista Imaging Tool v2.3 and stored in a database along with the measurements and possible remarks on the preservational state and/or shape and ornamentation of the chitinozoan vesicles. The database compiled for each of the samples can be found in attachments 9.3.1 to 9.3.18.

4.1.1.5 DATABASE AND STATISTICAL PROCESSING All chitinozoan photographs, measurements and remarks were stored in a database created with Filemaker Pro© 9.0v3. From the database, the species genus, species name, species number, lithology and sample number were exported to a Microsoft Office Excel worksheet for the compilation of an abundance matrix on genus level (Fig. 4.4) and species level (Fig. 4.5). Additional matrices were created by consecutively eliminating values lower or equal than 2 and lower or equal than 3 for both the genus level and species level.

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FIGURE 4. 4: ABUNDANCE MATRIX GENUS LEVEL. LIMESTONE SAMPLES ARE LIGHT GREY. SHALE SAMPLES ARE DARK GREY. SAMPLES WITH ONLY ONE OCCURRENCE OF A GENUS ARE MARKED SLIGHTLY DARKER, SAMPLES WITH ONLY TWO OCCURRENCES OF A GENUS ARE MARKED DARKER. TOTAL COUNT GIVEN ON THE RIGHT SIDE. NOTE: SAMPLES ARE IN ALPHABETIC ORDER, NOT STRATIGRAPHIC ORDER.

The constructed abundance matrices were then inserted in the Paleontological Statistics software Past 2.17c (Hammer, Harper & Ryan, 2001) for further data analysis which allowed to compare the samples. Analyses carried out on the dataset were inter alia a hierarchical cluster analysis, which can be used in identifying groups and subgroups in multivariate dataset, based on a given similarity measure. In this study, ‘paired group’ was chosen as algorithm along with both Euclidean and Morisita as similarity measure, the latter being frequently used in paleoenvironmental studies as it effectively normalizes for absolute abundances which makes this similarity comparatively insensitive to sample size. The paired- group algorithm (Unweighted pair-group average or UPGMA) works with clusters that are joined based on the average distance between all members in the two groups. This is preferential for ecological data (Hammer, 2006). All cluster analyses were bootstrapped 1000 times to assess their robustness. The result of a cluster analysis is displayed in a dendrogram, which branches dichotomously with each branching point representing a joining event during the agglomeration (the combining of clusters of similar objects). The branching point is drawn at a level corresponding to the similarity between the joined objects - indicating the degree of separation between clusters. It has to be noted that cluster analysis is no statistical procedure and that no significance level can be assigned to the formed clusters. It is, however, a useful method for data exploration - which was the aim of using this method in study. Various cluster analyses have been run on different abundance matrices on both the genus level and species level. They will be presented together with the results further on.

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FIGURE 4. 5: ABUNDANCE MATRIX ON SPECIES LEVEL. LIMESTONE LIGHT GREY. SHALE DARK GREY. ABUNDANCE <2 ONE SHADE DARKER, ABUNDANCE <3 TWO SHADES DARKER.

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A second analytical method that was applied was a correspondence analysis, a preferential method for projecting a multivariate dataset into two dimensions for visualizing trends and groupings. In paleoecology and paleobiogeography, correspondence analyses are frequently used for the ordination of both samples and taxa on the same plot (Hammer, 2006). Correspondence analysis tries to position both samples and taxa in the same space while still maintaining a correspondence between the two, somewhat similar to principal component analysis. Correspondence analysis uses indirect ordination, ordering samples solely by their taxonomic content, without a priori knowledge of their environmental position. This makes correspondence analysis ideal for ordination of samples and taxa that respond in a unimodal way (i.e. favoring a certain range) to an environmental gradient - being represented by the lithology in this study. If one would use principal component analysis, this would impose a linear response. Each axis of the correspondence analysis has its own Eigenvalue, indicating the degree to which ordination of data along the axis is successful. The first axis explains the most variation in a data set, the second axis the second most variation and so on - the higher the order of the axis, the less variation will be explained. In this study, a biplot is used to visually represent the results of the correspondence analysis. In practice, the primary underlying environmental gradient frequently influences both the first and second ordination axis, causing the samples and taxa to line up in an arch when displayed in a biplot. This can be avoided by using a detrended correspondence analysis, forcing the arch in a straighter shape while spreading the data more evenly (conform to normalization of data). Additionally, sample numbers and taxa are displayed on the biplot, facilitating the interpretation of which factors are at play.

When drawing conclusions from one analytical method (i.e. correspondence analysis in this study), it can be expedient to employ a second analytical method that has the ability to confirm or disprove the conclusions obtained by using the former method. For this, Principal Coordinate Analysis (PCO) has been utilized. PCO allows to project multivariate data into two dimensions, allowing the visualization of any possible trends or groups in the investigated dataset. The graphical representation of the data is based on the criterion that the Euclidean distances in this low-dimensional (2D) space should reflect the original distances in multidimensional space. Put in a more simplistic way; two similar data points should end up in the vicinity of each other in a PCO plot. Another advantage of PCO is that the type of similarity index or measure can be chosen freely; allowing the use of similarity indices that are optimized for ecological analyses, such as the Morisita similarity index (Hammer, 2006) that is relatively insensitive for sampling size because it normalizes for total abundances. One drawback for PCO is that the associated Eigenvalues can be negative, therefore compromising the validity of the analysis because the calculated distances should possess positive values.

All of the aforementioned methods are analytical and aimed at exploration of the dataset. It is desirable, however, to ascribe some statistical significance to the observations made based on the above methods. A non-parametric test for establishing statistically significant differences between two (or more) groups of multivariate data commonly used in ecology and biogeography is ANOSIM (Analysis of similarities). This type of test is frequently used to test for differences in faunal composition between the facies or regions (Harper, 2006). ANOSIM works by comparing within-group and across-group distances, and a test statistic R is based on the idea that the within-group distance should be significantly lower than the

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 across-group distance. The distance measure most frequently used in ANOSIM is the one of Bray-Curtis, which does not include absences in the data and does not normalize for abundances. This makes the distance measure very sensitive for abundant taxa (Harper, 2006). To assess if abundant taxa influence the outcome too much, another ANOSIM is run with the Morisita similarity index which normalizes absolute abundances. The significance of the R test is estimated by permutating samples across groups (Harper, 2006), displayed by the p(same)-value (p(same) = estimated probability of equality). The R test is considered significant if p(same) is lower than 0.05 (Harper, 2006).

4.2 GEOCHEMISTRY The geochemical analyses were carried out at the Friedrich-Alexander Universität Erlangen-Nürnberg, Germany, and therefore only a short explanation of the methods used will be given below. Ultimately, the goal of the geochemical analyses is to characterize both the limestone and shale by their chemical composition and compare them to one another to assess if the alternations were formed due to environmental changes or by differential diagenesis.

The analyses were performed by the use of XRF (X-Ray Fluorescence), an analytical method based on the characteristic secondary (fluorescent) X-Rays that are emitted from a material and its constituents after it has been bombarded with high-energy X-Rays. The signature of each element is based on the wavelength that is associated with the difference in energy when an electron returns to its pre-excited state, which is characteristic for each element. The intensity of each recorded wavelength is directly related to the abundance of the element in the sample.

4.2.1 THEORETICAL BACKGROUND During the early diagenetic stages, carbonate dissolution can take place in the marl layers and precipitates in the pore spaces of the developing limestone layers, thereby occluding the primary porosity of the limestone and hindering their compaction. This diagenetic redistribution can cause relative enrichment of insolubles in the marl and a dilution of insolubles in limestone in the diagenetically mature succession. Systematic differences in the ratios of insolubles can reveal primary compositional differences between marl and limestone layers because not every insoluble behaves similarly during diagenesis (Egger et al., 2013).

However, there is no possibility in unequivocally proving the absence of a primary depositional signal and therefore a diagenetic origin. This is because a primary signal could be concealed by parameters altered by diagenetic processes. This part of the geochemical analysis, based on the methodology of Munnecke et al. (2001, 2004), only allows differentiating between calcareous rhythmites with strong depositional variations and rhythmites without, but is unable to point out if the rhythmites are the result of rhythmic diagenesis of homogeneous precursor sediment or of the diagenetic enhancement of an underlying sedimentary rhythmic signal (Munnecke et al., 2001). Consequently, this geochemical analysis can only provide positive proof for an environmentally induced signal by distinct differences in

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 the lithologies; in contrast, the lack of an environmental signal does not allow discriminating between either of the two possibilities for the formation of calcareous rhythmites. Or as a common maxim states - absence of evidence is not evidence of absence. This explains our multi-proxy approach - a lack of environmental signal in multiple proxies would provide a stronger case for non-environmental, diagenetic steering of the facies alternations.

In addition, based on Munnecke et al. (2001, 2004), a possible composition of the hypothetical precursor sediment can be, assuming all soluble carbonate is derived from the diagenetically more unstable aragonite. This also allows tracing a diagenetic origin by assessing if the reconstructed precursor sediment composition meets realistic values. A full explanation for the employed methodology can be found in Munnecke et al. (2001) but the general outline is provided below. Research on the development of calcareous rhythmites from precursor sediments reveals that the carbonate cement required for the cementation of the limestone beds is derived by the preferential dissolution of aragonite in the adjacent marl layers. Consequently, the maximum available amount of carbonate cement equals the initial amount of aragonite present in the precursor sediments, which consists of a mixture of aragonitic, calcitic and terrigenous materials. From this observation, mathematical relationships have been conveyed between the mineralogical composition of the precursor sediments on the one hand, and the carbonate contents and thickness ratios of the alternations on the other. Using these relationships, it is possible to calculate the mineralogical composition of the precursor sediments of calcareous rhythmites and possibly reveal an environmental signal that otherwise might have remained hidden by the differential diagenesis affecting the diagenetically matured sections. As a prerequisite, the thickness of the studied limestone and marl layers needs to be known, along with the carbonate content of the respective layers and porosity measures for each individual limestone and marl layer. The calculations can be made with the help of an online applet developed by Munnecke & Westphal (2004) (http://www.gzn.uni- erlangen.de/~munnecke/lma/).

Another approach on the geochemical analysis is made by the investigation of the major and trace element ratios throughout the different lithologies. Three groups can be recognized among the elements that allow characterization of the rhythmites, of which the first group comprises oxides and elements which are bound to clay minerals, such as TiO2, K2O and Rb. The second group comprises elements and oxides that can be included in the lattices of both calcite and clay minerals, such as MgO,

Fe2O3 and Zn. The last group contains Sr and MnO which are normally bound to carbonates. By investigating the dispersal of these elements throughout the succession, information can be obtained whether or not diagenesis occurred and if it did; to what extent it altered the deposits.

4.2.2 SELECTED SAMPLES In total, 42 samples have been sent to the Friedrich-Alexander Universität Erlangen-Nürnberg, forming a continuous section from the top of the Solvang Formation downwards at the Bygdøy locality, representing the upper 4.55 meter of the Solvang Formation and consisting of 16 couplets of a limestone and marl layer.

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5. RESULTS For this study, a total of 46 samples were examined; 44 samples from the Solvang Formation at the Bygdøy locality and 2 samples from the Solvang Formation on Nakkholmen. The information obtained from analyzing and processing these samples will be given below.

5.1 PALYNOLOGICAL RESULTS A total of 18 samples have been dissolved, picked and investigated with the use of SEM. This resulted in 1636 photographs of specimens being made of which 1612 were positively identified as Chitinozoa. Out of those 1612 Chitinozoa, 1535 specimens were successfully identified to the genus level and 1209 specimens down to species level. In total, 11 different genera have been encountered and 16 different species were identified with confidence. Of those 11 different genera, 58.96% belongs to the genus Belonechitina and 20.05% to the genus Cyathochitina, therefore 79.48% or nearly 4 out of 5 of all encountered specimens belongs to only 2 of the 11 encountered genera. Moreover, on a species level, just over half of the specimens (52.18%) belong either to B. hirsuta complex (28.60%) or to B. robusta (23.58 %).

Below, three figures are given. Figure 5.1 is depicting the full range of genera encountered in all the samples, followed by Figure 5.2 depicting all the genera encountered in the continuous section and as a third, Figure 5.3 depicts all the species that occur in quantities greater than 5% in any sample for the continuous section.

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FIGURE 5. 1:GENUS BAR WITH TABLE DISPLAYING THE COMPOSITION OF EACH SAMPLE BY MEANS OF CHITIINOZOA GENERA. SAMPLES ARE ORDERED ACCORDING TO STRATIGRAPHIC HEIGHT. NOTE THE HIGH ABUNDANCE OF SPINACHITINA FOR THE FOUR LOWERMOST SAMPLES AND AN INCREASE IN CYATHOCHITINA IN THE SAMPLES IN THE MIDDLE.

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FIGURE 5. 2: BAR CHART ON GENUS LEVEL DISPLAYING THE COMPOSITION OF THE SAMPLES FOR THE STRATIGRAPHIC CONTINUOUS SECTION NEAR THE TOP OF THE SOLVANG FORMATION.

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FIGURE 5. 3: PERCENTAGES OF ALL SPECIES THAT OCCUR IN QUANTITIES LARGER THAN 5% IN ANY OF THE SAMPLES OF THE CONTINUOUS SECTION. HORIZONTAL ACCES RANGES FROM 0-75% FOR ALL GRAPHS. FIGURED CREATED WITH C2 DATA ANALYSIS V.1.7.2.

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5.1.1.1 DESCRIPTION OF ENCOUNTERED SPECIES Classification of the encountered species followed the proposals of Paris et al. (1999). All the species encountered in this study have been well described earlier, so only a short description along with eventual remarks is given as well as the publication in which the species are formally defined. If a certain genus only occurs once, its description in this study is omitted because of the low taxonomic value that can be assigned to this kind of singular identifications.

Incertae sedis group Chitinozoa Eisenack 1931; Order Operculatifera Eisenack 1972; Family Desmochitinidae Eisenack 1931; Subfamily Desmochitininae Paris 1981; Genus Calpichitina Wilson & Hedlund 1964;

Calpichitina spp. Fig. 5.4a Description: Lenticular vesicle, well developed colarette. Wall ornamentation not well enough preserved in encountered specimens. Dimensions: L: 45; Dp: 80-100; Dc: 45-70 Occurrence: WSA-12-083B (shale) Count: 3

Incertae sedis group Chitinozoa Eisenack 1931; Order Operculatifera Eisenack 1972; Family Desmochitinidae Eisenack 1931; Subfamily Desmochitininae Paris 1981; Genus Desmochitina Eisenack 1931;

Desmochitina erinacea Eisenack, 1931 Fig. 5.4b Description: ovoid to subspherical Desmochitina species with rough wall that is clearly ornamented with cones and tiny spines, evenly over the vesicle wall. Dimensions: L: 70-85; Dp: 50-70; Dc: 30-50 Occurrence: WSA-12-094 (limestone) Count: 8

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Desmochitina juglandiformis Laufeld, 1967 Fig. 5.4c Description: walnut shaped Desmochitina species with wrinkled wall and conspicuously protruding basal structure. Dimensions: L: 130; Dp: 70; Dc: / Occurrence: WSA-12-093 (shale) Count: 1

Desmochitina minor Eisenack, 1931 Fig. 5.4d Description: subspherical Desmochitina species with wide range in chamber morphology and glabrous surface. Dimensions: L: 80-110; Dp: 60-90; Dc: 30-60 Occurrence: WSA-12-094 (limestone), WSA-12-093 (shale), WSA-12-091 (shale), WSA-12-087 (shale), WSA-12-085A (shale), WSA-12-084 (limestone), WSA-12-083B (shale) Count: 32

Desmochitina ovulum Eisenack, 1962a Fig. 5.4e Description: the flaring colarette is clearly differentiated from the amphora-like to ovoid chamber. Smooth vesicle wall. Maximum width within upper half of the chamber. Dimensions: L: 100-120; Dp: 60-70; Dc: 30-40 Occurrence: WSA-12-101B (shale), WSA-12-093 (shale), WSA-12-091 (shale), WSA-12-085A (shale), WSA-12-084 (limestone), WSA-12-083B (shale) Count: 17

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Incertae sedis group Chitinozoa Eisenack 1931; Order Prosomatifera Eisenack 1972; Family Conochitinidae Eisenack 1931; Subfamily Conochitininae Paris 1981; Genus Conochitina Eisenack 1931; Fig. 5.4f Description: Only determinations on genus-level were made for Conochitina, grouped as Conochitina indet. due to the poor preservational state of the material. Conochitina indet. was used for Conochitinidae with a conical to claviform chamber and a glabrous wall. It is possible that some species belonging to the genus Belonechitina are included because their characteristic spines were eroded from the vesicle wall. Dimensions: no dimensions will be given due to the overall poor preservational state of the material. Occurrence: WSA-12-102 (limestone), WSA-12-101B (shale), WSA-12-091 (shale), WSA-12-090 (limestone), WSA-12-087 (shale), WSA-12-086 (limestone), WSA-12-085A (shale), WSA-12-085B (shale), WSA-12-084 (limestone), WSA-12-083B (shale), WSA-12-082 (limestone), WSA-12-031 (limestone), WSA-12-030B (shale) Count: 65

Incertae sedis group Chitinozoa Eisenack 1931; Order Prosomatifera Eisenack 1972; Family Conochitinidae Eisenack 1931; Subfamily Conochitininae Paris 1981; Genus Euconochitina Taugourdeau 1966 Fig. 5.4g Description: Conochitinidae with a conical chamber and glabrous wall. No ornamented vesicle walls. Euconochitina indet. was used for all specimens belonging to Euconochitina due to the poor preservational state of the studied material. Dimensions: no dimensions will be given due to the overall poor preservational state of the material. Occurrence: WSA-12-092 (limestone), WSA-12-090 (limestone), WSA-12-085A (shale), WSA-12-085B (shale), WSA-12-082 (limestone), WSA-12-031 (limestone) Count: 10

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Incertae sedis group Chitinozoa Eisenack 1931; Order Prosomatifera Eisenack 1972; Family Conochitinidae Eisenack 1931; Subfamily Conochitininae Paris 1981; Genus Rhabdochitina Eisenack 1931;

Rhabdochitina magna Eisenack, 1931 Fig. 5.4h Description: very long, slender Rhabdochitina species with a rounded basal margin and a rounded to flat base. Not as slender as Rhabdochitina gracilis, distinction made on Dp values. Dimensions: L: 290-400; Dp: 110-140; Dc: 100 Occurrence: WSA-12-087 (shale), WSA-12-085B (shale) Count: 4

Incertae sedis group Chitinozoa Eisenack 1931; Order Prosomatifera Eisenack 1972; Family Conochitinidae Eisenack 1931; Subfamily Belonechitininae Paris 1981; Genus Belonechitina Jansonius 1964;

Belonechitina hirsuta complex Vandenbroucke, 2008a, p.75-76 Fig. 5.4i Description: this group comprises small cylindro-conical to trapezoid, spiny specimens that can otherwise be attributed to the trapezoid to pear-shaped Belonchitina hirsuta or more ovoid Angochitina communis on grounds of their morphological similarity and the subsequent confusion between the two. In the studied samples, the two extremes shapes were identified but also a continuum of intermediate species was encountered. Therefore, it was opted to group all the small spiny chitinozoans into the Belonechitina hirsuta complex. The spines can vary between simple spines to lambda-form spines to multipodal. Also see Vandenbroucke et al. (in subm.) for similar practice. Dimensions: L: 80-130; Dp: 40-80; Dc: 15-45 Occurrence: in all the studied beds with the exception of WSA-12-101B (shale) Count: 439

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Belonechitina micracantha Eisenack, 1931 Fig. 5.4j Description: Belonechitina species with distinct cylindrical neck and conical chamber shape, shoulders are absent with straight flanks. The base is slightly convex. Only one specimen is identified in this study. Dimensions: L: 200; Dp: 90; Dc: 50 Occurrence: WSA-12-091 (shale) Count: 1

Belonechitina robusta Eisenack, 1959 Fig. 5.4k Description: Belonechitina species with cylindrical neck and conical chamber. The flexure is inconspicuous and the flanks are straight to slightly convex. Small lambda to multipodal spines are evenly distributed over the vesicle wall and are best developed towards the base of the vesicle. Dimensions: L: 150-350; Dp: 70-130; Dc: 50-90 Occurrence: in all the studied samples with the exception of WSA-12-102 (shale) but that may be due to the low number of specimens encountered in the particular sample Count: 362

Belonechitina wesenbergensis brevis Eisenack, 1972 Fig. 5.4l Description: Belonechitina species with cylindrical neck and conical chamber. The shoulders are well developed and the mostly simple, stout spines are well-developed and evenly distributed over the vesicle wall. The flanks are straight to slightly curved, the base is flat to slightly round. Dimensions: L: 100-160; Dp: 65-85; Dc: 25-50 Occurrence: WSA-12-091 (shale), WSA-12-089 (shale), WSA-12-085B (shale), WSA-12-084 (limestone), WSA-12-083B (shale) Count: 28

Incertae sedis group Chitinozoa Eisenack 1931; Order Prosomatifera Eisenack 1972; Family Conochitinidae Eisenack 1931; Subfamily Belonechitininae Paris 1981; Genus Hercochitina Jansonius 1964; Fig. 5.4m Description: Only two specimens of Hercochitina were identified to genus level in this study. They are Conochitinidae with a conical chamber and distinct crests consisting of vertically aligned spines. The flexure is visible and the shoulders are well developed. Dimensions: L: 145-160; Dp: 70-75; Dc: 40 Occurrence: WSA-12-030B (shale) Count: 2

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Incertae sedis group Chitinozoa Eisenack 1931; Order Prosomatifera Eisenack 1972; Family Conochitinidae Eisenack 1931; Subfamily Spinachitininae Paris 1981; Genus Spinachitina Schallreuter 1963;

Spinachitina bulmani Jansonius, 1964 Fig. 5.4n Description: Cylindro-conical Spinachitina species with a distinct neck and chamber and obvious (though gentle) flexure. Flanks are swollen, shoulders vague and the basal margin is sharp and bears a few large to numerous small spines, occasionally with multiple bases. The neck flares slightly towards the aperture. Dimensions: L: 130-220; Dp: 50-70; Dc: 30-50 Occurrence: WSA-12-101B (shale), WSA-12-031 (limestone), WSA-12-030B (limestone) Count: 52

Spinachitina multiradiata Eisenack, 1959 Fig. 5.4o Description: Cylindro-conical Spinachitina species with gentle flexure and straight to typically concave flanks, a sharp basal margin decorated with long basal processes which are mostly simple but can be complex. The neck flares towards the aperture and the crown might be ornamented with small, simple spines. In comparison to S. bulmani, is Spinachitina multiradiata is more Erlenmeyer-shaped. Dimensions: L: 120-170; Dp: 60-70; Dc: 25-50 Occurrence: WSA-12-101B (shale), WSA-12-031 (limestone), WSA-12-030B (limestone) Count: 39

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Incertae sedis group Chitinozoa Eisenack 1931; Order Prosomatifera Eisenack 1972; Family Lagenochitinidae Eisenack 1931; Subfamily Cyathochitininae Paris 1981; Genus Cyathochitina Eisenack 1955b;

Cyathochitina campanulaeformis Eisenack, 1931 Fig. 5.4p Description: Cyathochitina species with a cylindrical neck, strongly developed flexure and a swollen to conical bell-shaped chamber and a rather short carina. Obvious striations on the neck. Dimensions: L: 200-250; Dp: 130-160; Dc: 30-90 Occurrence: WSA-12-102 (limestone), WSA-12-101B (shale), WSA-12-090 (limestone), WSA-12-089 (shale), WSA-12-088 (limestone), WSA-12-087 (shale), WSA-12-086 (limestone), WSA-12-085A (shale), WSA-12-085B (shale), WSA-12-084 (limestone), WSA-12-082 (limestone), WSA-12-031 (limestone), WSA-12-030B (shale) Count: 70

Cyathochitina kuckersiana Eisenack, 1962 Fig. 5.4q Description: Cyathochitina species with a cylindrical neck, strongly developed flexure and a swollen to conical bell-shaped chamber and a longer, flaring carina. The vesicle bears obvious striations on the neck. On the well preserved samples, a central basal mark can be retrieved. Discussion: Easy to be confounded with C. campanulaeformis, as the distinction is made on length of the carina. Because of this, uncertainties identification can exist in samples with poor preservation. Dimensions: L: 130-300(395); Dp: 110-160; Dc: 30-90 Occurrence: WSA-12-102 (limestone), WSA-12-101B (shale), WSA-12-090 (limestone), WSA-12-089 (shale), WSA-12-088 (limestone), WSA-12-087 (shale), WSA-12-086 (limestone), WSA-12-085A (shale), WSA-12-085B (shale), WSA-12-084 (limestone), WSA-12-082 (limestone), WSA-12-031 (limestone), WSA-12-030B (shale) Count: 98

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Incertae sedis group Chitinozoa Eisenack 1931; Order Prosomatifera Eisenack 1972; Family Lagenochitinidae Eisenack 1931; Subfamily Ancyrochitininae Paris 1981; Genus Ancyrochitina Eisenack 1955a;

Ancyrochitina onniensis Jenkins, 1967 Fig. 5.4r Description: Small, cylindro-conical Ancyrochitina species with a flaring neck (if preserved), sharp flexure, a wide conical chamber with sharp basal margin which bears a few larger and branching spines whereas the wall is ornamented with numerous small, simple to lambda spines evenly distributed. The flanks are straight to slightly swollen. Dimensions: L: 95-130; Dp: 60-90; Dc: 30-50 Occurrence: WSA-12-091 (shale), WSA-12-089 (shale), WSA-12-085A (shale), WSA-12-085B (shale), WSA-12-084 (limestone), WSA-12-083B (shale), WSA-12-082 (limestone) Count: 47

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FIGURE 5. 4: RECOGNIZED SPECIES. A: CALPICHITINA SPP. (A.1 45-100-70; A.2 /-80-50). B: DESMOCHITINA ERINACEA (B.1 75-58-40; B.2 80-70- 50). C: DESMOCHITINA JUGLANDIFORMIS (130-70-/). D: DESMOCHITINA MINOR (D.1 120-80-40; D.2 90-70-50). E: DESMOCHITINA OVULUM (E.1 13-70-30; E.2 120-70-30). F: CONOCHITINA INDET. (F.1 180-90-70; F.2 200-70-60; F.3 160-60-30). G: EUCONOCHITINA INDET. (135-45-20). H: RHABDOCHITINA MAGNA (400-110-100). I: BELONECHITINA HIRSUTA COMPLEX (I.1 110-55-25; I.2 100-50-25). MEASUREMENTS L-DP-DC, ALL IN MICROMETER.

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FIGURE 5. 4 CONTINUED: RECOGNIZED SPECIES. I: BELONECHITINA HIRSUTA COMPLEX (CONTINUED) (I.3 100-70-30; I.4 100-70-35; I.5 100-65- 30; I.6 110-65-40; I.7 90-60-20). J: BELONECHITINA MICRACANTHA (200-90-50). K: BELONECHITINA ROBUSTA (K.1 180-100-70; K.2 230-120- 90; K.3 130-60-40; K.4 290-110-60; K.5 320-90-50). L: BELONECHITINA WESENBERGENSIS BREVIS (L.1 160-65-35; L.2 110-60-30; L.3 120-80-40; L.4 100-80-40). M: HERCOCHITINA INDET. (145-70-40). MEASUREMENTS L-DP-DC, ALL IN MICROMETER.

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FIGURE 5. 4 CONTINUED: RECOGNIZED SPECIES. N: SPINACHITINA BULMANI (N.1 140-40-35; N.2 180-60-30). O: SPINACHITINA MULTIRADIATA (O.1 140-70-40; O.2 170-60-35; O.3 120-70-35). P: CYATHOCHITINA MULTIRADIATA (P.1 200-150-70; P.2 240-150-70; P.3 220- 130-30). Q: CYATHOCHITINA KUCKERSIANA (Q.1 280-200-50; Q.2 230-130-50; Q.3 220-145-50). R: ANCYROCHITINA ONNIENSIS (R.1 95-75-30; R.2 100-70-30; R.3 130-90-50; R.4 100-70-40; R.5 120-80-40). MEASUREMENTS IN L-DP-DC, ALL IN MICROMETER.

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5.2 GEOCHEMICAL RESULTS In total, 19 samples out of 42 have been analyzed thus far as part of parallel student projects at the Friedrich-Alexander Universität Erlangen-Nürnberg (see attachment 9.4). Of those 19 samples, 17 samples were limestone and only 2 samples were shale. Values are reported in percentages for the main element oxides (SiO2, TiO2, Al2O3, Fe2O3, MnO, MgO, CaO, Na2O, K2O, P2O5) and the total Loss On Ignition (LOI), being indicative for the total amount of organic carbon. Additionally, the concentration of 12 trace elements (Ba, Cr, Ga, Nb, Ni, Pb, Rb, Sr, Th, V, Y, Zn, Zr) has been measured and their values are reported in parts per million (ppm). Starting from the main element oxides, the percentages of each pure element have been calculated. A next step was the calculation of the ratios of the pure main elements by normalization against Al, since aluminium oxides do not dissolve at the usual pH and redox-conditions that occur in deposits when diagenetically altered.

The calculation of a hypothetical precursor sediment (Munnecke & Westphal, 2004) has not been carried out by the lack of information on the carbonate content of the respective layers. This information will hopefully be available within the next months.

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6. DISCUSSION In this chapter, an exploration to the meaning of the presented results is made, i.e. what information can or cannot be derived. This is done separately for the palynological results and the geochemical results. Afterwards, a holistic interpretation of the palynological and geochemical results is reported.

6.1 IN-DEPTH EVALUATION OF THE PALYNOLOGICAL RESULTS A first exploration of the data can be accomplished by performing a cluster analysis including all the samples. The parameters were set with the Morisita similarity index (see Chapter 4.1.1.5) and a bootstrap value of 1000 in order to assess the robustness of the constructed dendrogram (Fig. 6.1).

FIGURE 6. 1: FULL SPECIES CLUSTER ANALYSIS WITH MORISITA SIMILARITY INDEX AND 1000 BOOTSTRAP. BASAL SAMPLES ARE SET APART FROM THE CONTINUOUS SECTION 100% OF THE ITERATIONS AND SHOW LOW SIMILARITY WITH THE OTHER SAMPLES. NO CLEAR DISTINCTION BETWEEN LITHOLOGIES IN THE CONTINUOUS SECTION. SHALE DARK FONT, LIMESTONE LIGHT FONT.

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For 100% of the performed cluster analyses, the entire group from the basal part of the formation cluster separately from the continuous section higher up in the succession at the lowermost branch of the dendrogram. This indicates that the basal four samples form a separate group that has a different species composition from the rest of the analyzed samples, at roughly ~ 25% similarity. In order to confirm this grouping as suggested by the cluster analysis, a correlation analysis was carried out. In the initial correlation scatter plot, the samples all lined up in an arch form, so therefore a detrended correlation analyses was made (Fig. 6.2) which confirms the observations made earlier, judging by the dendrogram.

FIGURE 6. 2: DETRENDED CORRESPODENCE ANALYSIS. THE SHALE OF THE CONTINUOUS SECTION IS OUTLINED BY THE DARK GREY HULL, THE LIMESTONE BY THE LIGHT GREY HULL. THE BASAL SAMPLES PLOT TO THE RIGHT, SEPARATELY FROM THE CONTINUOUS SECTION. SHALE: DARK GREY HULL; LIMESTONE: LIGHT GREY HULL.

The detrended correspondence analysis scatter plot depicts two clusters on the left side (i.e. the shale samples of the continuous section enclosed by the dark-shaded grey hull and the limestone enclosed by the light-shaded grey hull) whereas the samples that constitute the ‘basal’ group of samples are isolated

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 on the right-hand side of the plot, clearly distinct from those from the stratigraphically-higher continuous section.

Based on these two analyses, it is shown that withholding the base samples from the continuous section for further analyses is justified. Moreover, the basal samples would introduce a stratigraphical signal to the dataset that might blur the ecological influence that we seek to explore. The four basal samples have purposely only been analyzed to assess distinct changes in faunal composition between the topmost part of the formation and its bottommost layers. Their inclusion in the same figures and graphical representations of the rest of the dataset would give a deceitful impression of sudden changes in the chitinozoan assemblages whereas it solely indicates an actual biostratigraphic difference in assemblage composition, while nothing can be concluded on the nature of that change - i.e. whether it is a sudden or gradual change remains to be revealed by the investigation of further samples. Samples WSA-12-030B, WSA-12-031 represent the base of the Solvang Formation at Bygdøy and WSA-12-101B and WSA-12-102 at Nakkholmen.

Statistical verification to the previous methods of assessment can be achieved with the ANOSIM analysis, illustrating whether the base samples truly form a statistical significant different group. The test was run on the complete abundance matrix using the Bray-Curtis (see Chapter 4.1.1.5) distance measure. The basal samples were grouped and compared to all the other samples, regardless of lithology. The R test, part of the ANOSIM analysis, returns a value of 0.9516 and thereby indicates that the groups are very different. The low p(same)-value of 0.0004 indicates that this result of the R test is significant and thus it can be concluded that the mixed bottom samples are distinctively different from the stratigraphically higher samples. Individual comparisons of the base shale samples compared to the top shale samples return R = 0.987 and p(same) = 0.0288 whereas limestone base samples to limestone samples return R = 0.9026 and p(same) = 0.0259. So even within the same lithology, the base samples remain significantly different from the others.

6.1.1 SPECIES LEVEL For the remainder of the analyses, the 4 lowermost samples or so-called ‘basal’ samples (WSA-12-030B, WSA-12-030, WSA-12-101B, WSA-12-102) are removed from the dataset to limit the influence of non- ecological signals. Subsequently, a cluster analysis, correspondence analysis, PCO analysis and ANOSIM analysis performed on the species using the abundance matrix that only includes the samples from the continuous section. No clear separation in assemblages is observed during the analyses, but a few trends can be discerned, based on the abundances of certain species.

6.1.1.1 CLUSTER ANALYSIS The first exploration of the dataset, consisting of all the different species encountered (Fig. 4.5, with the species occurring in the base samples omitted from the dataset) was through construction of 2 dendrograms by applying a cluster analysis. The first cluster analysis was run by using the Morisita similarity measure (see Chapter 4.1.1.5), thereby eliminating any influence of sample size, and a

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 bootstrap value of 1000 (Fig. 6.3) whereas the second cluster analysis was run with the Euclidean similarity measure (see Chapter 4.1.1.5), the simplest distance measure, and a bootstrap value of 1000 (Fig. 6.4). Both the similarity measures are applied to the abundance matrix, allowing for mutual verification of the constructed dendrograms.

FIGURE 6. 3: FULL SPECIES CLUSTER ANALYSIS WITHOUT BASAL SAMPLES, USING EUCLIDEAN SIMILARITY INDEX AND 1000 BOOTSTRAP. ALL LITHOLOGIES ARE MIXED. FOUR SAMPLES GROUP SEPARATELY ON THE RIGHT HAND SIDE FOR ALL ITERATIONS. FURTHER GROUPINGS OF LOWER CONFIDENCE AS INDICATED BY BOOTSTRAP VALUES. SHALE DARK FONT, LIMESTONE LIGHT FONT.

At a first glance, it is already clear that there is no obvious division between the two lithologies based on their palynological content since the main branches constructed in the dendrogram comprise both lithologies. Furthermore, the bootstrap indicates that the robustness of the constructed dendrogram (or reproducibility) is rather low so not too much confidence can be held in the analysis.

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FIGURE 6. 4: FULL SPECIES CLUSTER ANALYSIS WITHOUT BASAL SAMPLES USING MORISITA SIMILARITY INDEX AND 1000 BOOTSTRAP. ALL LITHOLOGIES ARE MIXED. FOUR SAMPLES GROUP SEPARATELY ON THE RIGHT HAND SIDE FOR ALL ITERATIONS. FURTHER GROUPINGS OF LOWER CONFIDENCE AS INDICATED BY BOOTSTRAP VALUES. SHALE DARK FONT, LIMESTONE LIGHT FONT.

Analyzing the same dataset utilizing the same analytical method but with slightly different parameters (i.e. different similarity measure) can help to confirm the conclusions made from Figure 6.3. This is achieved by performing a cluster analysis with the Euclidean similarity measure, being the basic algorithm for expressing distances between data points. The dendrogram in Figure 6.4 confirms the previous observation that the lithologies show no distinct grouping. One feature present in both dendrograms is the separation of WSA-12-088, WSA-12-091, WSA-12-092 and WSA-12-094 from the rest of the samples, and as indicated by the bootstrap values, this is achieved for every iteration of the analysis. An explanation for this separate grouping can be the high content in B. hirsuta complex that characterizes each of the samples (> ~50 %.), though the separation is not related to the sampled facies.

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6.1.1.2 CORRESPONDENCE ANALYSIS Dataset examination by cluster analysis indicated that there is no distinct or obvious grouping of the dataset in two, nicely separated clusters of lithologies by their chitinozoan content at the species level. Validation for this observation is given by applying a methodologically more robust correspondence analysis; allowing the discovery of groupings and underlying environmental gradients in a multivariate dataset consisting of taxonomic counts in a number of samples. The results of the analysis are displayed in Figure 6.5.

FIGURE 6. 5: FULL SPECIES CORRESPONDENCE ANALYSIS. THE INFLUENCE OF CERTAIN SPECIES IS NOTICEABLE (CALPICHITINA SPP., C. KUCKERSIANA, A. ONNIENSIS, B. WESENBERGENSIS BREVIS). THE TWO LITHOLOGIES STILL SHOW OVERLAP. SHALE: DARK GREY HULL; LIMESTONE: LIGHT GREY HULL.

At a first glance, the two lithologies (shale is the dark-grey hull, limestone is the light-grey hull) form more or less separated groupings, yet there is still considerable overlap between the groups. The horizontal (first) axis accounts for only 33% of the variation in the sample, the vertical axis for 18% of the total variation. These are rather low Eigenvalues, and given that the 3rd and 4th Eigenvalue still account

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 for ~10% of the variation each, it is concluded that there is no strong, singular influence at play in the present dataset on species level. When studied in detail, the influence of the difference species can be assessed. For example, Calpichitina spp. plots all the way up on top in the graphic, thereby confirming is does not occur in any sample except WSA-12-083B, which also plots the closest to Calpichitina spp. Similarly, the influence of D. erinacea can be noted on WSA-12-094, as it only occurs in that sample. The highest abundances for the genus Ancyrochitina are found in shale samples, which is made clear by their central position amongst the shale samples. Several species with low abundances (mostly in WSA-12- 091 and WSA-12-093; e.g. B. microacantha, D. juglandiformis, Lagenochitina sp., D. minor) plot central on the diagram, and pull WSA-12-091 and WSA-12-093 towards them. A final remark might be that C. kuckersiana and C. campanulaeformis plot low on the diagram, in the vicinity of the limestone samples.

6.1.1.3 PCO ANALYSIS The correspondence analysis on species level seems to suggest that to some extent, the two lithologies might be characterized by distinct assemblages of chitinozoan. An additional PCO analysis assesses the validity of the correspondence analysis; by verifying if the same conclusions can be drawn using the former method. To this effect, a PCO analysis was run with the Morisita similarity index (Fig. 6.6; see Chapter 4.1.1.5). The two hulls by the lithologies mostly collide, except for the limestone samples WSA- 12-088, WSA-12-092 and WSA-12-094 that group to the left exterior of the shale-hull. These samples (together with WSA-12-091, which is also positioned quite to the left) were the same samples that grouped separately on the cluster analysis because of their high B. hirsuta complex content. The first two axes represent in total 74.5% of the variation in the sample (61% and 13.5% respectively), which is significantly more than the correspondence analysis. This PCO analysis also indicates that the suggested grouping by the correspondence analysis might not be as significant as it appears to be.

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FIGURE 6. 6: FULL SPECIES PCO ANALYSIS WITHOUT BASAL SAMPLES. THE TWO LITHOLOGIES LARGELY COLLIDE. THE FOUR LEFTMOST SAMPLES STAND OUT BECAUSE OF THEIR HIGH CONTENT IN B. HIRSUTA COMPLEX. SHALE: DARK GREY HULL; LIMESTONE: LIGHT GREY HULL.

6.1.1.4 ANOSIM All the previous methods applied were aimed at a first examination of the data. Thus far, the data revealed no clear grouping of the species into assemblages characteristic for each lithology. As the assessments were made on visual basis, no truly objective significance can be given to them. Therefore, an ANOSIM test was performed on the dataset to assess if there is a significant difference between the two faunas of the two lithologies. The first ANOSIM was run on the abundance matrix without the base samples and with the Bray-Curtis distance measure. The limestone samples are compared with the shale samples. The R test gave a value of 0.07191 indicating that there is not much difference between distances measured within-group (44.24) and across-group (47.51). The value of 0.2109 for p(same) is above the significance level (i.e. >0.05). Hence, ANOSIM confirms that the assemblages of the lithologies are not significantly different.

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6.1.2 GENUS LEVEL The analyses of the abundance matrix on species level did not reveal a direct, clear link between the lithofacies and assemblage thus far. To verify if an environmental signal is present within the different lithofacies on a genus level, the same array of analyses has been applied to the genus-level abundance matrix.

6.1.2.1 CLUSTER ANALYSIS The cluster analysis was run using the Morisita similarity index with a bootstrap value of 1000 (Fig. 6.7). Again, a clear grouping of the lithofacies is absent from the unmodified genus abundance data matrix. Samples WSA-12-082, WSA-12-084, WSA-12-085B, WSA-12-086 and WSA-12-087 group separately from the rest with high confidence, which can be explained by their high content in Cyathochitina.

FIGURE 6. 7: FULL GENUS CLUSTER ANALYSIS WITHOUT THE BASAL SAMPLES USING MORISITA SIMILARITY INDEX AND 1000 BOOTSTRAP. THE 5 SAMPLES ON THE RIGHT BRANCH DISTINCT THEMSELVES BY THEIR CYATHOCHITINA CONTENT. NO OBVIOUS GROUPING IN LITHOLOGY. SHALE DARK FONT, LIMESTONE LIGHT FONT.

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6.1.2.2 CORRESPONDENCE ANALYSIS Similar to the method followed on species level, a correspondence analysis was made to visually represent any possible groupings or environmental factors in the dataset (Fig. 6.8). This graph confirms that no clear distinction can be made between the lithofacies as their hulls collide. Similar to the correspondence analysis, the influence of Calpichitina and the positioning of WSA-12-083B are notable. Furthermore, Belonechitina, which occurs in all samples in significant to large quantities, plots central on the graph, in near vicinity of most samples. The genus Ancyrochitina lies in the center of the hull described by the shale samples, reflecting the higher occurrence of Ancyrochitina in the shale samples. The first axis represents 45 % of the total variation, the second axis 24 %. A PCO analysis was run to confirm the findings by the correspondence analysis and returned very similar results.

FIGURE 6. 8: FULL GENUS CORRESPONDENCE ANALYSIS WITHOUT BASAL SAMPLES. THE LITHOLOGIES LARGELY COLLIDE. INFLUENCE OF CALPICHITINA, BELONECHITINA AND CYATHOCHITINA ON THE DISTRIBUTION OF THE SAMPLES. SHALE: DARK GREY HULL; LIMESTONE: LIGHT GREY HULL.

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6.1.2.3 ANOSIM The ANOSIM test returned an R value of -0.004373 which indicates that there is more difference within the lithologies than between them. Or; there is more variance within the genus assemblage than there is between the lithologies. The value of 0.395 for p(same) also confirms that there is no significant grouping possible.

6.1.3 PUSHED ANALYSES TO TEASE OUT MAXIMUM ENVIRONMENTAL SIGNAL Chitinozoa are not generally characterized by a strong differentiation between shelfal and non-shelfal species. In fact, their highest diversity is at high latitude in less distal environments (Grahn et al., 2011; Vandenbroucke et al., 2010). Instead, their spatial distribution seems to a large degree controlled by SST (Vandenbroucke et al. 2010 a, b), also hypothesized to change at our sampling site if the rhythmites represent cyclic paleoclimatic signals. However, our site might be out of spatial reach of the migrating water masses, and a not-unimportant portion of the fauna consists of generalist species (Vandenbroucke et al. 2010 a, b). Therefore, we want to test if there are sub-assemblages of selected species that do seem to preferentially occur in different biofacies (or lithofacies), i.e. either the odd species that is controlled by depositional depth, or the more sensitive ones that react to less important SST changes. We thus test if the abundance matrices constructed in this study, might reveal an environmental signal hidden among the ‘background noise’ created by the bulk of the species. For this, the percentages for each species for both the limestone and the shale have been calculated. Next step was filtering on percentages higher than one standard deviation (68.2%) and for low occurrences (species must occur in more than 1 sample), therefore maximizing the discrepancy between the two lithologies. As a result, only 7 species were withheld in a new abundance matrix (Fig. 6.9) that was subsequently analyzed by performing a cluster analysis (Fig. 6.10) with the Morisita similarity index and a bootstrap value of 1000.

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FIGURE 6. 9: MODIFIED SPECIES ABUNDANCE MATRIX. SPECIES WITHHELD HAVE AN OCCURRENCE OF MORE THAN 68% IN EITHER SHALE OR LIMESTONE AND OCCUR IN MORE THAN ONE SAMPLE. BASAL SAMPLES NOT INCLUDED.

The grouping seems more confined to the lithologies. The first obvious feature is the separate grouping of WSA-12-093 and WSA-12-094, which can be linked to the total absence of Cyathochitina kuckersiana in both samples. The rest of the samples groups in two lithological distinct groups, safe for WSA-12-087 that is grouped with limestone samples. The grouping has a rather low robustness (bootstrap value indicates 27% reproducibility). When looked at in detail, WSA-12-087 shares more characteristics with the limestone by the absence of A. onniensis and B. wesenbergensis brevis and a higher abundance for C. kuckersiana than compared to the other shale samples. This observation is confirmed by a correspondence analysis (Fig. 6.11) showing that WSA-12-093 and WSA-12-094 group separately on the top left and WSA-12-087 plotting amidst the limestone samples. Since WSA-12-082 and WSA-12-084 contained few A. onniensis, they plot closer to the shale samples than the other limestone samples. The scatter plot describes 68.5% of the total variation by its first two axes (40.5% and 28% respectively).

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FIGURE 6. 10: SELECTED SPECIES CLUSTER ANALYSIS WITHOUT BASAL SAMPLES USING MORISITA SIMILARITY INDEX AND 1000 BOOTSTRAP. SHALE DARK FONT, LIMESTONE LIGHT FONT

A closer look at Figure 6.9 reveals that WSA-12-094 only contained D. minor after the species selection, the rest of the sample comprised of 76.3% B. hirsuta complex, an additional 10.5% B. robusta and a few specimens of D. erinacea, all of which were not withheld in the modified species matrix. If WSA-12-094 is omitted from the correspondence analysis, two distinct groupings show on the scatter plot (Fig. 6.12).

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FIGURE 6. 11: SELECTED SPECIES CORRESPONDENCE ANALYSIS WITHOUT BASAL SAMPLES. WSA-12-093 AND WSA-12-094 PLOT SEPARATELY BECAUSE OF THEIR D. MINOR CONTENT. INFLUENCE OF C. KUCKERSIANA ON THE ONE HAND AND B. WESENBERGENSIS BREVIS AND A. ONNIENSIS ON THE OTHER ARE CLEAR. SHALE: DARK GREY HULL; LIMESTONE: LIGHT GREY HULL.

All limestone samples clusters tightly together around C. kuckersiana, whereas the shale samples form a broader group. Although all the limestone samples contain other species than C. kuckersiana, they only contain few of them. Vice versa, most of the shale samples also contain C. kuckersiana, albeit in low abundances. This suggests that C. kuckersiana might be a species that preferentially thrives in environmental settings that are beneficial for the deposition of limestone at this site and therefore serves an index species for identifying biofacies - i.e. it would be a shelfal species. This is also mentioned by Nestor (1998) who states that C. kuckersiana preferentially occurs in shallow-water conditions whereas C. campanulaeformis is intermediate between shallow water and deep water conditions - which are characterized by C. calix. An alternate explanation could be that C. kuckersiana preferentially

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Wout Salenbien The nature of calcareous rhythmites in the Ordovician of Oslo 2012 - 2013 occurs (i.e. has higher abundances) in more temperate/colder waters, which are supposed to be more prevalent in the study area during times of glaciation when global sea levels are lower since massive amounts of water are stored in expansive ice caps, causing colder water bodies to move towards the equator and the tropics to be reduced in size. In this regard, C. calix could be seen as indicative for warmer sea surface temperatures and being more abundant in the study area when sea levels are high during interglacials. C. campanulaeformis would have a higher tolerance for sea surface temperature fluctuations and would therefore be controlled to a lesser extent these fluctuations. Interestingly, Vandenbroucke et al. (2010a, Fig. 9), admittedly in another time slice and keeping in mind that little is known about the drift of chitinozoans’ tolerance to SST through time, identified C. kuckersiana as a polar to subtropical generalist species with the lowest latitudinal occurrences at ~30°S. The only (doubtful) occurrence of B. wesenbergensis brevis in Vandenbroucke et al. (2010a) is at low latitudes (25°S).

FIGURE 6. 12: SELECTED SPECIES CORRESPONDENCE ANALYSIS WITHOUT BASAL SAMPLES AND WITHOUT WSA-12-094. ALL LIMESTONE SAMPLES CLUSTER TIGHTLY AROUND C. KUCKERSIANA WHEREAS THE BULK OF THE SHALE SAMPLES IS CONCENTRATED NEAR A. ONNIENSIS AND B. WESENBERGENSIS BREVIS. NOTE THE LOWER AFFINITY OF WSA-12-087 AND WSA-12-093 WITH THE REST OF THE SHALE SAMPLES. SHALE: DARK GREY HULL; LIMESTONE: LIGHT GREY HULL.

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It should be stressed that care has to be taken with either of these conclusions. The distinction between C. kuckersiana and C. campanulaeformis - the other Cyathochitina species occurring in relatively high abundances - is based on the length of the carina and therefore relying strongly on the preservational state of the chitinozoans in the samples. Whether the lithological alternations were formed by a true alternation in depositional settings or formed by differential diagenesis from an uniform precursor calcareous mud, the shale layers are more susceptible to deformation in later events (e.g. the Caledonian orogeny, Permian rifting) than are the limestone layers. This can cause the fragile carina to erode, therefore tampering the possible identification of C. kuckersiana or any other species. When looked at all Cyathochitina occurring in the samples and the distribution between limestone and shale, it can be noted that for C. campanulaeformis, the count is more or less similar for both lithologies with 26 counts in limestone samples to 18 counts in shale samples - being more or less even for both lithologies. For C. kuckersiana, the difference is, as noted before, more pronounced with 83 specimens in limestone and only 15 in shale samples. But when looked at C. indet., which was used to group all the remaining Cyathochitina that did not clearly belong to either of the two other recognized species, 66 specimens are counted in limestone compared to 76 counts for the shale. This indicates that for the shale samples, preservational conditions were slightly worse. And even more, some C. campanulaeformis belonging to either limestone or shale could have belonged to C. kuckersiana but could have been identified as C. campanulaeformis if the distinct longer carina was eroded. Or, some species could have belonged to C. calix, which is distinct from C. campanulaeformis and C. kuckersiana by its greater length. But rarely a Cyathochitina vesicle is preserved well-enough to display its full length, therefore making it difficult to recognize C. calix in the poorly-preserved samples, only one specimen was identified in the shale sample WSA-12-030B. If in the correspondence analysis is run without C. kuckersiana, no grouping can be seen anymore and all the limestone samples plot within the shale samples. Paris & Verniers (2005) remark that C. campanulaeformis displays a more or less global distribution.

Another remark can be made on the occurrence of A. onniensis. As mentioned before, the species seems to preferentially occur in shale samples. In fact, 4 out of 7 shale samples group closely by the species on the correspondence analysis and the two limestone samples that contain A. onniensis are also drawn towards the species on the plot. But still, 83% of the occurrences of A. onniensis is in shale lithology, which is about the same lithological discrepancy that can be noted for C. kuckersiana, albeit for opposite lithologies. The very same holds true for B. wesenbergensis brevis, of which 85% of the encountered specimens is to be found in shale deposits, as to the remaining 15% in the limestone. From this, A. onniensis and B. wesenbergensis brevis could be viewed as indicative for deeper conditions - perhaps non-shelf. Or, alternatively, indicators of warmer sea surface temperatures during interglacials that are characterized by higher sea levels.

From this discussion, it can be concluded that - on a species level and solely based on this restricted study - no real indication of a significant different assemblages as a whole can be identified for each of the lithologies, but some species can potentially serve as facies indicator or SST-indicator. This is partly due to the poor preservational state of the palynomorphs, introducing uncertainties in species identification. An ANOSIM test was run to assess if the difference between the groups is significant. The R test reports a value of 0.3565 - indicating that there is some difference, but no a clear grouping

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(within-group distance 37.26 compared to across-group distance 53.49) from the lithologies. The p(same) value of 0.0151 indicates that the grouping is statistically significant.

If an analogues filtering methodology is used for the genera abundance matrix, only 2 genera remain with enough discrepancy between limestone and shale. This does not allow for a correspondence analysis, so the filtering criteria were slightly lowered to 66% (or ~2/3). If a correspondence analysis (Fig. 6.13) is made from the matrix, no clear distinction can be seen between the lithologies, as they are largely colliding. Both WSA-12-090 and WSA-12-092 only contain 1 Euconochitina specimen and they plot far from the other samples. The influence of the genus Ancyrochitina on the shale samples is noticeable near the bottom of the scatter plot, causing the shale to plot slightly apart. The high content of Desmochitina in WSA-12-094 and WSA-12-093 causes them to group close together. Although WSA- 12-085A has a similar content in Desmochitina, its substantial amount of Ancyrochitina causes the sample to plot between the two genera.

FIGURE 6. 13: SELECTED GENUS CORRESPONDENCE ANALYSIS WITHOUT BASAL SAMPLES. INFLUENCE OF EUCONOCHITINA ON WSA-12-090 AND WSA-12-092, WHO ONLY CONTAIN THIS GENUS AFTER SELECTION. THE BULK OF THE SAMPLES PLOTS AS A CENTRAL CLUSTER SINCE FEW ARE INFLUENCED BY ANY OF THE WITHHELD GENERA. SHALE: DARK GREY HULL; LIMESTONE: LIGHT GREY HULL.

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6.1.4 EXPLORING VARIABILITY WITHIN ONE LAYER During this study, and most of other palynological studies, it is always presumed that one sample is sufficient to represent the total variation within the layer that is being analyzed. While this might hold true for 5 to 10 centimeter thick samples, this could already be different for layers with a more substantial thickness, despite displaying no changes in lithology whatsoever. For this purpose, 2 samples from the 23 centimeter shale layer WSA-12-085 have been acquired during the fieldwork and subsequently have been dissolved and picked, with WSA-12-085A representing the lowermost 4 centimeter of the layer and WSA-12-085B coming from near the top. Both samples share some characteristics, as both contain A. onniensis and have very little C. kuckersiana, but some discrepancies arise as well. For the lower sample, WSA-12-085A, 47.58% of its content consists of species belonging to the Belonechitina genus compared to only 34.86% for WSA-12-085B. The same can be said for the Cyathochitina genus with 14.56% and 43.94% respectively. The genus Desmochitina makes up 16.51% of WSA-12-085A whereas none are retrieved in the other sample. The results of the comparison between the two subsamples have been plotted in Figure 6.14 on both species and genus level.

FIGURE 6. 14: LEFT: SPECIES LEVEL, LARGE DISCREPANCY FOR B. HIRSUTA COMPLEX, C. INDET., D. MINOR AND D. OVULUM. RIGHT: GENUS LEVEL, LARGE DISCREPANCY ON CYATHOCHITINA AND DESMOCHITINA. SCALE BARS FOR BOTH GRAPHS IN PERCENT. SHALE: DARK GREY HULL; LIMESTONE: LIGHT GREY HULL.

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This calls for caution when one is to rely on the presence or absence of a singular species to infer environmental information as it is clear from the above comparison that a sufficient coverage of the layer by the sample is indispensable. The same maxim as used before is applicable in this very situation; absence of evidence is not (always) evidence of absence.

6.1.5 CONCLUSIONS FOR THE PALYNOLOGICAL DATA The thorough analysis of the palynological data did not provide any clear evidence for different chitinozoan assemblages as a whole characterizing either the limestone or the shale, but did reveal 3 species that can be used in the Solvang Formation to infer deeper or shallower depositional settings (i.e. shelfal break or deeper versus shelf environment). A stratigraphical influence is noticeable when the four lowermost samples are included for the analysis; the presence of Spinachitina in those samples clearly sets them apart from the continuous section near the top of the Solvang Formation. As no samples in between are processed, it is unclear where the Spinachitina leaves/enters the assemblage, so no further conclusions can be drawn. When the continuous section is filtered to display maximum discrepancy between the lithologies, it still fails to show any distinct grouping on a genus-level. For the species level, a grouping can be inferred between the two lithologies, but due to the lower confidence in the identification of C. kuckersiana, this explanation should be treated with caution. When analyzed without C. kuckersiana on species level, the two lithologies largely collide again - showing that the clear grouping is dependent on one species and thus might rather be a preservational artifact. It must be noted that A. onniensis and B. wesenbergensis brevis show a preferential occurrence in shale deposits - which can be interpreted as facies indicator of deeper depositional settings or as indicative of the expansion of climate provinces with warmer sea surface temperatures. However, this does not show unambiguously on the correspondence analysis for the selected species when C. kuckersiana is left out (Fig. 6.15), as the limestone samples are completely within the field described by the shale samples and plot in the close vicinity of both A. onniensis and B. wesenbergensis brevis.

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FIGURE 6. 15: SELECTED SPECIES CORRESPONDENCE ANALYSIS WITHOUT BASAL SAMPLES, WSA-12-094 AND C. KUCKERSIANA. THE LIMESTONE PLOTS WITHIN THE VARIABILITY OF THE SHALE SAMPLES, YET STILL FORMS A TIGHT CLUSTER COMPARED TO THE SHALE SAMPLES. SHALE: DARK GREY HULL; LIMESTONE: LIGHT GREY HULL.

6.2 IMPLICATIONS FOR THE PALYNOLOGICAL DATA The fact that palynological data does not immediately allow for revealing differences in the assemblages between the lithologies - except for the previously discussed 4 species - has a whole range of possible explanations and consequences. It has been remarked before that Baltica moved from high latitudes towards lower latitudes during the Ordovician so that during the deposition of the Solvang Formation, the Oslo-Asker region would have been situated at c.35°S (Cocks & Torsvik, 2005) and would have possessed the intermediate biodiversity values characteristic for Baltica at that time. This would have had its influence on the composition of the chitinozoan fauna, being characterized by warmer water species compared to the Early Ordovician. In Vandenbroucke et al. (2010a), a reconstruction of the chitinozoan biotopes has been made for the Sandbian gracilis time slice (Fig. 1.11). A first order major boundary in biozones was reported to occur at c.35°S which formed the boundary between subtropical

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(towards the equator) and subpolar biotopes. It should be noted that there is a 5° uncertainty with these reconstructions. However, this latitude of 35°S would put the study area in a transitional zone intermediate between the two distinctly different biotopes. The boundaries of these biotopes are susceptible to change through time, as can be seen for the Hirnantian in Vandenbroucke et al. (2010b), so changes between the Sandbian and Katian may occur.

One could argue that changes in sea-level induced by waxing and waning of ice sheets has its influence on sea surface temperatures as global temperature conditions change, and that this very change in temperature can be recorded by the chitinozoan fauna. The studied Oslo-Asker region in Baltica, however, was located in on the verge between subtropical and subpolar chitinozoan biotopes where the sea surface temperature fluctuations between glacials and interglacials would have been of a smaller magnitude - similar as the reconstructed fluctuations during the Holocene. If glacio-eustatic sea level changes would occur, along with a change in climate, this can cause the transitional zone to show more characteristics from either of the adjacent biotopes. If this occurred, it is expected to show in the palynological assemblage.

When looked at the whole assemblage, no significant difference between the two lithologies can be deduced as confirmed by statistics. This can interpreted as that (1) there is no actual difference in assemblage and the rhythmites are the result of differential diagenesis or (2) the shift in biotopes by sea level change (and climate) is large, but not registered in the study area because it remains within the extremes of the changing biotopes (i.e. the chitinozoans are not sensitive enough to sea-level change) as sketched in Figure 6.16. A last possible scenario can be that (3) the changes in the latitudinal position of the climate belt and thus the biotope of the chitinozoans did not result in a change in deposition style at the study area (which is sea-level driven). This might have been because the study area did not leave the boundaries of the prevalent biotope/climate belt. A limited influence on the chitinozoans may be possible for the species that react more sensitively to environmental changes.

FIGURE 6. 16: SKETCH OF POSSIBLE EXPLANATION 2, SHOWING THE SHIFT WITHIN THE TRANSITIONAL BIOTOPE THAT DO NOT CAUSE A NEW BIOTOPE TO BE INSTALLED AT THE STUDY AREA.

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Another possible explanation, which is related to the previously mentioned one, could be that on this time scale chitinozoans are not that sensitive enough to register the changes in sea level that occur with ‘Milankovich’ sea level fluctuations. Hints et al. (2006) conclude in a study on chitinozoan biofacies in the Llandovery of western Estonia that the relative frequency of chitinozoan species not necessarily is related to changes in lithology and thus depositional or environmental settings and report the absence of change in the chitinozoan fauna across a lithologically sharp boundary with marked changes only occurring above the boundary.

However, if the chitinozoan abundance matrix is scrutinized for maximum discrepancy between lithologies, four species can be recognized that cause a statistically significant grouping of the lithologies into two assemblages. Safe for C. calix, which only has a singular occurrence, all species occur in both lithologies, but their abundances differ. If the view of the study area being located in the transitional zone between the subtropical and subpolar biotopes is held, the species preferentially occurring in the limestone can be interpreted as deposited under low sea levels that are likely to occur during glacials when large amounts of water are stored in ice caps. In contrast, the species occurring more abundantly in shale can be indicative of warmer SSTs, invading the sampling site while the higher sea levels, concomitant with the warmer climate of interglacials, allowed for the mudline to shift landwards and shale be deposited. The fact that this shift is only seen in a statistically significant way for some species can be explained by the intermediate-to-lower latitudinal location of Baltica, at which fluctuations in sea surface temperature were likely low in amplitude and may only have had a limited influence.

6.3 BIOSTRATIGRAPHY According to Bergstrӧm et al. (2011), the Solvang Formation is placed near the middle of the Katian at the top of the graptolite Dicranograptus clingani Biozone, being the topmost zone in the Baltoscandic Rakvere (E) stage or equivalent to the topmost part of the Katian 1 stage slice (Fig. 1.1) or time slice 5C according to Webby et al. (2004). Despite yielding a decent number of moderately preserved chitinozoans, the assemblage diversity is rather low and lacks stratigraphically important species that allow positioning the Solvang Formation more precisely in time. Species as C. campanulaeformis, C. kuckersiana, C. calix, D. minor, D. ovulum, B. micracantha, B. robusta and B. hirsuta complex are all species are all documented in the literature as having an extensive range through a large part of the Upper Ordovician.

None of the Chitinozoa zonal index species listed in Nõlvak et al. (2006) for the chitinozoan biozones and subzones in the Baltoscandian regional timescale occurs in the studied samples, so the stratigraphic position cannot be confirmed by using the biozonation as proposed by Nõlvak et al. (2006). However, some accessory species can be found that are able to suugest a certain biozone or at the very least a biochron if their temporal range is limited in the fossil record. In the four lowermost samples, derived from the Bygdøy locality and from Nakkholmen, an assemblage of Spinachitina multiradiata and Spinachitina bulmani is encountered, albeit in a poor preservational state so some caution concerning the identification should be taken. These species are known to occur concurrently with Spinachitina

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To conclude it can be remarked that the Solvang Formation ranges in chronostratigraphy from the Oandu Stage (suggested by the Spinachitina) into the Rakvere Stage as suggested by the occurrence of B. wesenbergensis brevis. A similar conclusion was reached by Grahn et al. (1994) based on chitinozoan biostratigraphy. This means that the Solvang Formation is situated on the boundary between S. cervicornis biozone and the F. spinifera biozone. This coincides largely with the D. clingani graptolite biozone, but the base of the formation is less clearly confined in time (Fig. 6.17).

It should be noted that Bergstrøm et al. (2011) confine the Solvang Formation to the upper part of the Rakvere Stage based on the occurrence of the KOPE isotopic excursion in the formation, which also allows correlation the Laurentia. Based on the apparent discrepancies between the chemostratigraphic and tentative biostratigraphic positioning of the Solvang Formation, future research is warranted and needed to elucidate the stratigraphic positioning of the Solvang Formation.

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FIGURE 6. 17: BIOSTRATIGRAPHIC SCHEME FROM PARIS ET AL. (2005) SHOWING THE BALTOSCANDIAN CHITINOZOAN BIOZONES FOR THE ORDOVICIAN AS WELL AS FOR OTHER PALEOCONTINENTS, ALLOWING QUICK CORRELATION.

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6.4 GEOCHEMISTRY The geochemical analysis are not completed yet, but although only 2 marl samples were analyzed in comparison to 17 limestone samples, this still allows for a preliminary interpretation of the geochemical data, providing that one keeps in mind the low number data that is readily available to draw our conclusions from.

As described with the presentation of the geochemical data earlier, all element ratio against aluminium have been calculated. The degree to which they correlate with Al allows them to subdivide them in a group with a very good positive correlation (Ti, Si, Zr, Fe). A group with good positive correlation with Al can be discerned that either shows some scatter or does not go through the zero-point of the diagram (K, Zn). A last group correlates negatively with Al (Mn, Sr).

The major and trace elements can be divided into groups according to the processes that mostly influence them. In that way, it is possible to recognize elements that are preferentially bound to clay minerals, such as Ti, Al, K, Rb and Si. Zr is believed to be controlled by the amount of coarser detrital input (Faure, 1986; Egger et al., 2013). The ratios are for Ti:Al, K:Al, Si:Al and Zr:Al all show an excellent positive correlation (R2 > 0.97) between the limestone and marl, even R2 > 0.99 for Si:Al and Ti:Al (Fig. 6.18). All these elements are believed to be quite stable during diagenesis, and the fact that they correlate for both lithologies indicates that there were no major changes in the detrital influx during deposition and the clay mineral composition was rather homogeneous during “limestone times” and “marl times”, although the amount of terrigenous input might have varied in between. Similar values and distrubtions have been found by Egger et al. (2013) in the Hovedøya Member of the stratigraphically higher Skogerholmen Formation. In all the graphics in Figure 6.18 a small subgroup within the limestone samples can be discerned (WSA-12-076, WSA-12-078, WSA-12-090) that plot somewhat away from the bulk of the limestones towards the shale samples. These could be interpreted as less mature limestone layers or could have been samples that simply were closer to the margin of the limestone layer.

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FIGURE 6. 18: RATIOS FOR SI, TI, K AND ZR ALL SHOW AN EXCELLENT TO GOOD CORRELATION WITH AL. ALL POINT TO LACK OF VARIATION IN CLAY MINERALS DURING DEPOSITION. SHALE SAMPLES PLOT TOP RIGHT IN EVERY DIAGRAM.

Another group of elements (Mg, Fe, Zn) can be identified by their preferential inclusion in the crystal lattice of calcite and clay minerals. The correlation between Al and these elements is less clear (R2(Fe) = 0.9812; R2(Mg) = 0.7693; R2(Zn) = 0.5259)). In case of the Zn:Al correlation, the significantly lower R2 can be related to 2 peak values in the limestone samples that cause the correlation to deteriorate (Fig. 6.19). A possible explanation could be the formation of sphalerite crystals within the sediment, but this needs to be further investigated. A peak value for Fe (5.56% Fe) has been removed from WSA-12-090 (limestone) as it dwarfed the percentages of Fe in all the other samples and may have been caused by a chunk of pyrite in the sample or some other oddity.

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FIGURE 6. 19: RATIOS FOR FE, MG AND ZN CORRELATED WITH AL. FE CORRELATES VERY GOOD, ZN AND MG SOMEWHAT LESS, CAUSES REMAIN COMPLICATED AND UNCLEAR. SHALE SAMPLES PLOT TOP RIGHT IN EVERY DIAGRAM.

A last group of elements are preferentially incorporated in the carbonates (Sr, Mn) and thus correlate negatively with aluminium. The correlations are somewhat less distinct (Fig. 6.20), with R2(Mn) = 0.5717 and R2(Sr) = 0.7177. The large variability may be caused by primary differences in the sediment, by the different diagenetic maturity of each limestone layer or by other numerous processes that can affect the distribution of these elements.

FIGURE 6. 20: RATIOS FOR MN AND SR WHICH BOTH CORRELATE NEGATIVELY WITH AL BECAUSE OF THEIR PREFERENTIAL UPTAKE IN CARBONATES. SHALE SAMPLES PLOT BOTTOM RIGHT IN EVERY DIAGRAM.

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Overall, enrichment in trace elements (Ba, Cr, Ga, Nb, Ni, Pb, Rb, Th, V, Y, Zr) for the marl layers can be observed with the exception of Sr-values. As the Ti:Al indicate more or less constant terrigenous influx throughout the succession, the lower content of trace elements in the limestone can be explained as dilution by the import of carbonate cement during early diagenesis.

If all the different ratios for the studied elements are taken into account, it can be concluded that the geochemical data points to a diagenetic origin of the calcareous rhythmites rather than an environmental origin as no major difference in clay mineral composition can be discerned throughout the limestone and clay units. However, this needs to be confirmed by the study of further samples. And, even if the signal remains indicative of diagenesis after the study of more samples, this indication by the geochemical data cannot serve as unambiguous evidence for an entirely diagenetic origin of the rhythmites because primary differences could have been developed by other parameters (such as porosity, content of organic material or the ratio of aragonitic to calcitic carbonate) that have been effaced by the subsequent diagenesis. Therefore the need arises to interpret the geochemical data while taking the palynological data into account.

6.5 AN INTEGRATED VIEW OF THE DATA At first, it seems as if the data and conclusions derived from the palynological analyses and the geochemistry contradict. The geochemical data clearly favors a diagenetic origin of the rhythmites, with the distribution of the investigated elements clearly correlating across the lithologies, indicating a common, or at the least very similar, original composition for both limestone and marl layers. Keeping in mind, of course, that absence of an environmental signal in this proxy record cannot unequivocally prove its absence. With the same constrain, the bulk of the palynological data, in the broad perspective of including all species, also rather favors the diagenesis-hypothesis since no distinct change in assemblages arises throughout the different lithologies and only one assemblage seems to exist throughout the studied section. There is no direct indication of an environmental control on the chitinozoan assemblage. However, if the chitinozoan abundance matrix is modified to display maximum discrepancy between lithologies, 4 species stand out that have significantly different abundances in the respective lithologies in this study and thus seem to favor one lithology (and the associated environmental/climatic parameters involved with the inferred separate deposition of the lithology) above the other. It has to be stressed that, according to other studies (Vandenbroucke et al.,2010b), C. kuckersiana is a generalist species for subpolar to subtropical biotopes and B. wesenbergensis brevis is assigned (however with low confidence) to low latitudes/tropic biotopes.

There are two potential ‘philosophies’ to interpret the data at hand;

1. ‘Weight of evidence’ reasoning: The weight of every source of evidence is considered in making an interpretation of the data. In this study, both the geochemical and the total palynological data favor for a diagenetic origin of the calcareous rhythmites whereas only a heavily filtered species matrix shows signs of environmental influences that have that are not free of uncertainties and pitfalls (e.g. C. kuckersiana is a non-specialist in other studies and its

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identification can be subject to facies-controlled preservational state). So the largest body of evidence dictates a diagenetic origin of the rhythmites from, presumably, rather homogeneous precursor sediment. 2. Compromise reasoning: Every piece of data is considered and combined in an integrated explanation that tries to account for the apparent absence of an environmental signal in the geochemistry and full palynological data, but also for the apparent signal in the stringently filtered dataset. It relies heavily on the environmental signal introduced by the four species in the modified abundance matrix. Such a possible ‘compromise’ is sketched below:

Short-term climate fluctuations might not have a large effect on the chitinozoans biotopes in the sampling site, as changes in sea surface temperature are limited near low-latitudes (Kucera, 2007); depending on the detailed paleogeography of our site it may or may not be impacted by the climate belt migrations Also, it is quite possible that the environmental changes are not large enough to install a different chitinozoan biotope at the study area (e.g. if the study area remains within the limits of the shifting bioprovince/climate belt). In addition, most chitinozoan species are known not to be very sensitive to water depth fluctuations, which nevertheless are the suggested driving force behind the registration of the environmental signal in the deposited sediment. Yet, some species - the ones most sensitive to SST, or that do relate to water depth - may show (subtle) differences in their abundances, driven by the environmental changes, and small differences in the deposited sediment may be present because of the slightly changing boundary conditions. This signal might be blurred by species that have a wider tolerance for changes in SSTs in case of the palynological data or, for the geochemical data, may have been effaced from the sedimentological record by later diagenesis. Thus it seems possible that the alternations are formed by diagenetic enhancement of (weak) differences in composition of the precursor sediment as suggested by the palynological content.

The combined data seems to support the idea that calcareous rhythmites reflect environmental changes, albeit not as contrasting and in a direct one-to-one manner as previously believed. It rather seems that small differences in more homogenous sediment are greatly enhanced by differential diagenesis which leads to the creation of the typical calcareous rhythmites. Despite this, their occurrence and origination still lend support to the interpretation as used in the Baltic sea level reconstruction by Nielsen (2004). However, further and more detailed research is very necessary to confirm and further extend the observations made in this study and to flesh out the evidence for this explanation.

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7. CONCLUSION This study was aimed at elucidating the origins of the calcareous rhythmites of the Solvang Formation in the Ordovician of the Oslo-Asker region and improving its (bio)stratigraphic significance by means of a study of the chitinozoan content and the geochemical characterization of the beds. By tracing the origins of the Solvang Formation, a better understanding of the development of the Early Paleozoic Icehouse can be obtained as the onset of the ice-cap formation can be allocated more precisely in time. The possible hypotheses on the formation of the Solvang Formation were either that the rhythmic alternations represent high-order environmental (possibly sea-level) fluctuations during a general glacio- eustatic lowstand, or, according to Munnecke et al. (2001), the calcareous rhythmites were formed via differential diagenesis from a much more homogeneous precursor sediment and therefore reflect little to no environmental (let alone climatological) signal.

In total, 18 samples have been studied for their chitinozoan content, resulting in the counting of 1612 Chitinozoa of which 1535 were well-enough preserved to be identified to genus-level and 1209 to species-level. An additional 42 samples, forming a ~4.5 meter continuous section from the top down of the Solvang Formation at Bygdøy, were sent to the Friedrich-Alexander Universität Erlangen-Nürnberg for geochemical analyses. This allowed characterizing part of the Solvang Formation both in terms of the chitinozoan content and geochemical character.

As of this writing, not all the samples intended for geochemical analysis have been processed so only preliminary results can be reported based on 17 limestone samples and only two shale samples. From the analysis of the major and trace elements and their ratios plotted against aluminium, it can be concluded that the two lithologies can be correlated to each other by their content as indicated by Ti, K, Si and Zr. Elements that preferably are incorporated in carbonates (Sr, Mn) correlate negatively with aluminum, which is to be expected. By this it can be concluded that the geochemical data favors a diagenetic origin of the rhythmites. Yet, it cannot be hold as conclusive evidence due to the nature of the method: only a clear facies-specific geochemical signal points to an environmental origin of the rhythmites. The absence of such a clear signal (as observed in the Solvang Formation) can reflect either that the environmental signal is not recorded by this proxy (moderate enhancements of lithological differences) or that this signal was never there, pointing to an origin of the facies that is complete due to differential diagenesis from a homogeneous precursor sediment. A more detailed report on the geochemical will hopefully be given by the end of June 2013 when more data on the samples will be available.

The palynological study of the Solvang Formation reveals no distinctly different assemblages on both the genus and species level, but does allow for recognizing 4 species that do show statistically significant changes in abundances for the two lithologies when the abundance matrix is modified to show maximum discrepancy between the limestone and marl.

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The data in this study can be interpreted in two ways.

(1.) This first one being that both the geochemical evidence and the bulk palynological evidence point to a diagenetic origin of the rhythmites, based on the absence of an obvious environmental signal in two independent proxies, known to register environmental signals in other times % places. Amberg (pers. comm.) and Egger et al. (2013), based on preliminary data, draw a similar conclusion for the calcareous rhythmites of the Hovedøya Member of the Skogerholmen Formation, higher up in the stratigraphy of the Oslo-Asker region

(2.) . When combining the findings of the palynological data and the geochemistry and attributing more significance to the four environmentally controlled species, a possible alternative explanation arises that is capable of explaining the conclusions reached by each of the analytical methods. Chitinozoans are controlled by sea surface temperature and, according to Vandenbroucke et al. (2010a,b), this allows to discern different biotopes similar to modern-day planktonic provinces. In the biotope reconstruction for the Sandbian gracilis-time slice, a first order major boundary between biotopes existed at ~30-35°S (with a 5° margin of error), roughly the latitude at which the Oslo-Asker region would have been located during the Katian, which divided the subtropical from the subpolar biotopes. At these paleolatitudes, a transitional biotope would have existed (Vandenbroucke et al., 2010a). In another reconstruction for much cooler Hirnantian times, this transitional biotope did not shift dramatically compared to its Sandbian position, to ~25-30°S (Vandenbroucke et al., 2010b). If, in addition, depositional settings would not have changed dramatically (though enough to produce primary differences) throughout the period in which the Solvang Formation was deposited and the changes in SST modest (analogous to the modern-day glacial-interglacial changes in SST), it can be envisaged that no large changes would be recorded in the chitinozoan assemblage at the study area. A minor influence on the abundance of some species is to be expected parallel along with some changes in the type of deposited sediment (e.g. small changes in carbonate content, organic material, porosity, etc.). For the sediment, these small differences can be effectively effaced by differential diagenesis - which occurred according to the geochemical data, and quite likely greatly enhanced the lithological differences. Only the palynological data is capable of suggesting a cyclic nature in this case, and 4 species are recognized to show statistically significant differences in abundance along with the changing lithologies.

Further research on the palynological characterization and geochemistry is needed to assess the reliability of the view proposed above due to uncertainties and a limited body of evidence. Other studies, using the same methodology as applied herein (Collart, 2013; Amberg, in prep.), may help verify or nullify the proposed interpretations. For now, it can be concluded that the interpretation of the calcareous, used by Nielsen (2004) in the construction of the Baltic sea level curve, can be supported by this study, albeit in a more nuanced view and relying on the validity of the environmental signal recorded by the four chitinozoans.

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9. LIST OF ATTACHMENTS

9.1 SAMPLING LISTS

9.1.1 HUK FORMATION 9.1.2 NAKKHOLMEN FORMATION - BYGDØY 9.1.3 SOLVANG FORMATION - BYGDØY 9.1.4 NAKKHOLMEN FORMATION - NAKKHOLMEN 9.1.5 SOLVANG FORMATION - NAKKHOLMEN 9.1.6 VENSTØP FORMATION - NAKKHOLMEN 9.1.7 GRIMSØYA/SKJERHOLMEN FORMATION - NAKKHOLMEN

9.2 LABORATORY LOGBOOK

9.3 CHITINOZOAN DATABASE 9.3.1 WSA-12-030B 9.3.2 WSA-12-031 9.3.3 WSA-12-082 9.3.4 WSA-12-083B 9.3.5 WSA-12-084 9.3.6 WSA-12-085A 9.3.7 WSA-12-085B 9.3.8 WSA-12-086 9.3.9 WSA-12-087 9.3.10 WSA-12-088 9.3.11 WSA-12-089 9.3.12 WSA-12-090 9.3.13 WSA-12-091 9.3.14 WSA-12-092 9.3.15 WSA-12-093 9.3.16 WSA-12-094 9.3.17 WSA-12-101B 9.3.18 WSA-12-102

9.4 GEOCHEMICAL ANALYSIS

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