The evolution of lunar U-Pb geochronology of Ca-phosphates and zircon using Secondary Mass Spectrometry Fiona Karen Thiessen Academic dissertation for the Degree of Doctor of Philosophy in Geology at Stockholm University to be publicly defended on Friday 9 November 2018 at 10.00 in De Geersalen, Geovetenskapens hus, Svante väg 14.

Abstract Planetary bodies in our Solar System, including the , were exposed to an intense bombardment between ~4.5-3.8 Ga, shaping their surfaces and leaving visible “footprints” in the form of large impact basins. The end of this period (~4.0-3.85 Ga), might have been marked by a cataclysmic increase in impacts, the so-called (LHB), although this remains highly contentious. Since destructive processes, such as tectonics or erosion, have destroyed ancient (> 3.0 Ga) impact structures on Earth, studies of the early Solar System are mainly restricted to lunar samples, because impact structures are much better preserved on the Moon. In this thesis, we have therefore analysed impact breccias from three landing sites (, 14, and 17) with the overall aim to gain a better understanding of the lunar impact history. This endeavour included comprehensive textural and petrological analyses of the breccias and grains of interest (i.e. Ca-phosphates and zircon), as well as obtaining precise U-Pb Secondary Ion Mass Spectrometry (SIMS) ages. The U-Pb ages of Ca-phosphates obtained are consistent with the age of the Imbrium impact at ~3925 Ma, whereas an older age of ~3930 Ma yielded by Ca-phosphates in an might be linked to the formation of the Serenitatis basin. Furthermore, an at ~3940 Ma was identified in zircon grains in Apollo 14 breccias, which is in agreement with older Ca-phosphate ages yielded in a previous study. The identification of three possible impact events within ~15 myr has important implications for the lunar bombardment history. However, there is a possibility that partial Pb loss from older grains during a relatively late event (e.g. Imbrium) might result in apparently older ages in Ca-phosphates. Incomplete resetting of the U-Pb system was recorded in zircon grains in an Apollo 12 breccia, leading to meaningless U-Pb ages which cannot be interpreted unambiguously as either magmatic or as impact events. Nevertheless, the U-Pb ages of several zircon grains occurring in lithic clasts in Apollo 14 breccias can plausibly be linked to magmatic activity, exhibiting several magmatic events between ~4286 Ma and ~4146 Ma. The data obtained in this thesis, together with previously published zircon and Ca-phosphate data, indicate several spikes in the magmatic and impact history during the first ~600 myr of lunar history. This study highlights the importance of combining high-precision age determination with thorough petrological and textural analyses in order to exclude meaningless ages and to interpret the impact and magmatic history of the Moon.

Keywords: Moon, impact breccias, Apollo missions, impact event, Late Heavy Bombardment, zircon, Ca-phosphates, U- Pb, SIMS, geochronology.

Stockholm 2018 http://urn.kb.se/resolve?urn=urn:nbn:se:su:diva-160018

ISBN 978-91-7797-396-6 ISBN 978-91-7797-397-3

Department of Geological Sciences

Stockholm University, 106 91 Stockholm

THE EVOLUTION OF LUNAR BRECCIAS

Fiona Karen Thiessen

The evolution of lunar breccias

U-Pb geochronology of Ca-phosphates and zircon using Secondary Ion Mass Spectrometry

Fiona Karen Thiessen ©Fiona Karen Thiessen, Stockholm University 2018

ISBN print 978-91-7797-396-6 ISBN PDF 978-91-7797-397-3

Cover: The Moon and the Apollo 12, 14, and 17 landing sites. The background image is from the Lunar Reconnaissane Orbiter Camera (NASA/GSFC/Arizona State University). : Picture taken by Claudia Bessa Görmarker

Printed in Sweden by Universitetsservice US-AB, Stockholm 2018 To my grandfather, Günter Adamski

Sammanfattning Solsystemets planetära kroppar, månen inkluderad, utsattes för ett intensivt asteroid-bombardemang för ca 4,5 till 3,8 miljarder år sedan, vilket formade deras yta och lämnade tydliga ”fotspår” i form av stora impakt-bassänger. Slutet av denna period (ca 4,0 till 3,85 miljarder år sedan) kan ha präglats av en katastrofal ökning av antalet impakter, detta kallas det sena tunga bombardemanget (eng: ”Late Heavy Bombardment (LHB))”, även om detta fortfarande är mycket omtvistat. Eftersom destruktiva processer, såsom tektonik och erosion, har förstört gamla (äldre än 3,0 miljarder år sedan) nedslagsstrukturer på jorden, är studier av det tidiga solsystemet huvudsakligen begränsade till prover ifrån månen, eftersom nedslagstrukturerna är mycket bättre bevarade på månen.

I denna avhandling har vi därför analyserat olika breccior ifrån tre olika Apollo landningsplatser (Apollo 12, 14 och 17) med det övergripande målet att få en bättre förståelse för månens nedslagshistoria. Detta projekt innefattade omfattande texturella och petrologiska analyser av brecciorna samt specifika mineral av särskilt intresse (dvs Ca- fosfater och zirkon), samt att erhålla U-Pb åldrar med hjälp av jonmikrosond (SIMS). U-Pb-åldrarna av Ca-fosfater överensstämmer med åldern av Imbrium-impakten vid ca 3925 Ma, medan en äldre ålder på ca 3930 Ma ifrån Ca- fosfater ur Apollo 17-breccior kan vara knuten till bildandet av Serenitatis-bassängen. Det identifierades även ett nedslag vid ca 3940 Ma i zirkonkristaller från Apollo 14 breccior, vilket överensstämmer med en äldre Ca-fosfat- ålder som redovisats i en tidigare studie. Tre möjliga nedslag inom 15 miljoner år har identifierats, vilket kan leda till viktig information om månens bombardemangshistoria.

Det finns emellertid risk att det skett partiell Pb-förlust från äldre kristaller under en i sammanhanget relativt sen händelse (t.ex. Imbrium-impakten), vilket kan ge skenbart äldre åldrar i Ca-fosfater. Ofullständig nollställning av U- Pb-systemet noterades i vissa zirkonkristaller i en Apollo 12-breccia, vilket i sammanhanget leder till åldrar som inte kan tolkas entydigt som antingen magmatiska eller som ålder för impakten. Likväl kan U-Pb åldrarna hos flera zirkonkristaller som förekommer i bergartsfragment i Apollo 14-breccior sannolikt kopplas till magmatisk aktivitet, och de uppvisar flera faser av magmatism mellan ca 4286 Ma och ca 4146 Ma. Datan som presenteras i denna avhandling tillsammans med tidigare publicerad zirkon- och Ca-fosfatdata tyder på en mer dramatisk historia under månens första 600 miljoner år. Denna studie exemplifierar vikten av att kombinera precisa åldersbestämningar med noggranna petrologiska och texturella analyser för att utesluta intetsägande åldrar och på så vis förstå månens magmatiska historia samt dess nedslagshistoria.

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Zusammenfassung Zwischen ~4.5–3.8 Ga waren die planetaren Körper in unserem Sonnensystem – einschließlich des Mondes – einem intensiven Bombardement von Asteroiden ausgesetzt, welches sichtbare Spuren in Form von großen Einschlagsbecken hinterlassen hat. Ende dieser Periode (~4.0–3.85 Ga) könnte durch einen katastrophalen Anstieg an Asteroideneinschlägen gekennzeichnet worden sein, und wird auch als „Großes Bombardement“ bezeichnet. Diese Theorie bleibt aber weiterhin sehr umstritten. Da alte Impaktstrukturen (> 3 Ga) auf der Erde durch Prozesse wie Tektonik oder Erosion zerstört wurden und auf dem Mond um einiges besser erhalten sind, sind Studien des frühen Sonnensystems hauptsächlich auf Mondproben beschränkt.

In dieser Doktorarbeit haben wir daher verschiedene Impaktbrekzien von drei Apollo Landestellen (Apollo 12, 14 und 17) analysiert, mit dem allgemeinen Ziel ein besseres Verständnis der Einschlagsgeschichte des Mondes zu erlangen. Dieses Vorhaben umfasste eine strukturelle und petrologische Analyse der Brekzien und Körner (Calciumphosphate und Zirkone), sowie eine präzise Uran-Blei Sekundär-Ionen-Massenspektrometrie (SIMS) Altersbestimmung. Die erhaltenen U-Pb der Ca-Phosphate stimmen überein mit dem Alter des Imbrium Impakts vor ~3925 Ma, wobei ein älteres Alter von ~3930 Ma für Ca-Phosphaten in Apollo 17 Brekzien erhalten wurde, welches eventuell mit der Bildung des Serenitatis Einschlagbeckens in Verbindung gebracht werden kann. Des Weiteren wurde ein Alter von ~3940 Ma in Zirkonen in Apollo 14 Brekzien identifiziert, welches übereinstimmt mit älteren Altern von Ca-Phosphaten aus einer früheren Studie. Die Identifizierung von möglicherweise drei Impakten innerhalb einer Zeitspanne von 15 Millionen Jahre hat wichtige Implikationen für die Einschlagsgeschichte des Mondes.

Es besteht jedoch die Möglichkeit, dass das U-Pb-Zerfallssystem von älteren Ca-Phosphaten während eines späteren Events (z.B. der Imbrium-Impakt) teilweise zurückgesetzt wurde und diese daher nur scheinbar älter sind. Ein unvollständiges Zurücksetzen der radiometrischen Uhr wurde in Zirkonen in Apollo 12 Brekzien festgestellt, welches zu bedeutungslosen U-Pb Altern führt. Diese können weder eindeutig als Einschlagsalter noch als magmatische Kristallisationsalter interpretiert werden. Die U-Pb Alter mehrerer Zirkone, die in Gesteinsfragmenten in Apollo 14 Brekzien vorkommen, konnten dennoch magmatischer Aktivität zugeordnet werden und mehrere magmatische Ereignisse zwischen ~4286 Ma und ~4146 Ma wurden identifiziert. Die Daten, die in dieser Doktorarbeit erhalten wurden, zusammen mit bereits veröffentlichten Ca-Phosphat und Zirkon-Daten, deuten auf mehrere Höhepunkte der magmatischen und Einschlagsgeschichte des Mondes während der ersten 600 Millionen Jahre hin. Diese Studie hebt hervor, wie wichtig es ist hochpräzise Altersdatierung mit einer gründlichen petrologischen und strukturellen Analyse zu verbinden, um bedeutungslose Alter auszuschließen und um die Einschlagsgeschichte und die magmatische Geschichte des Mondes zu interpretieren.

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List of papers and authors´ contributions This thesis is composed of four papers listed below. Paper I is published in & Planetary Science, Paper II is published in Geochimica et Cosmochimica Acta, and Paper III is submitted to Meteoritics & Planetary Science.

I. Thiessen F., Nemchin A. A., Snape J. F., Whitehouse M. J. and Bellucci J. J. (2017). Impact history of the Apollo 17 landing site revealed by U‐Pb SIMS ages. Meteoritics & Planetary Science, 52(4), 584–611.

II. Thiessen F., Nemchin A. A., Snape J. F., Bellucci J. J. and Whitehouse M. J. (2018). Apollo 12 breccia 12013: Impact-induced partial Pb loss in zircon and its implications for lunar geochronology. Geochimica et Cosmochimica Acta, 230, 94–111.

III. Thiessen F., Nemchin A. A., Snape J. F. and Whitehouse M. J. (submitted) U-Pb SIMS ages of Apollo 14 zircon: Deconvolving distinct magmatic episodes.

Paper I: Thiessen F. carried out SEM imaging/mapping and SIMS U-Pb isotope analyses, interpreted the data, and wrote the manuscript. Whitehouse M. J., Nemchin A. A., Snape J. F., and Bellucci J. J. provided support for data acquisition, interpretation of the data, and manuscript corrections. The sample thin sections were prepared at NASA Johnson Space Center.

Paper II: Thiessen F. carried out acquisition of the data (SEM and SIMS), participated in the data interpretation, and wrote the manuscript. Nemchin A. A. helped acquiring and interpreting the data and corrected the manuscript. Whitehouse M. J., Snape J. F., and Bellucci J. J. provided support for data acquisition and manuscript corrections. The sample thin sections were prepared at NASA Johnson Space Center.

Paper III: Thiessen F. participated in SEM imaging/mapping and SIMS U-Pb analyses, interpretation of the data and wrote the manuscript. Snape J. F. acquired the SEM data, helped with SIMS analyses, and the correction of the manuscript. Nemchin A. A. and Whitehouse M. J. provided support with data acquisition, interpretation, and manuscript corrections. The sample thin sections were prepared at NASA Johnson Space Center.

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The following manuscripts are not part of this thesis:

I. Thiessen F., Besse S., Staid M. I. and Hiesinger H. (2014). Mapping units in as observed with the Moon Mineralogy Mapper (M³). Planetary and Space Science, 104, 244–252.

II. Snape J. F., Nemchin A. A., Grange M. L., Bellucci J. J., Thiessen F. and Whitehouse M. J. (2016). Phosphate ages in Apollo 14 breccias: Resolving multiple impact events with high precision U–Pb SIMS analyses. Geochimica et Cosmochimica Acta, 174, 13–29.

III. Merle R. E., Nemchin A. A., Whitehouse M. J., Pidgeon R. T., Grange M. L., Snape J. F. and Thiessen F. (2017). Origin and transportation history of lunar breccia 14311. Meteoritics & Planetary Science, 52(5), 842–858.

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Contents Sammanfattning ...... 1 Zusammenfassung ...... 2 List of papers and authors´ contributions ...... 3 1. Introduction ...... 7 1.1 Thesis aim ...... 7 1.2 Early impact cratering (~4.4–3.8 Ga) and the proposed LHB ...... 7 1.2.1 Origin of the LHB ...... 9 1.2.2 Recent studies and impact ages ...... 10 2. Geological processes ...... 11 2.1 Impact cratering ...... 11 2.2 Magmatic evolution and major lunar provinces ...... 14 2.3 Lunar minerals and rocks ...... 15 3. Geological setting ...... 17 3.1 Apollo 12 ...... 18 3.2 Apollo 14 ...... 19 3.3 Apollo 17 ...... 20 4. Methods ...... 22 4.1 Impact Breccias ...... 22 4.1.1 Ca-Phosphates ...... 22

4.1.2 Zircon (ZrSiO4) ...... 23 4.2 U- Pb geochronology ...... 25 4.2.1 Wetherill concordia diagram ...... 27 4.2.2 Tera-Wasserburg diagram ...... 27 4.3 Initial lunar Pb ...... 28 4.4 Partial resetting ...... 29 5. Analytical methods ...... 30 5.1 Scanning microscopy (SEM) ...... 30 5.1.2 Cathodoluminescence (CL) imaging ...... 31 5.2 Secondary ion mass spectrometry (SIMS) ...... 31 6. Summary of manuscripts ...... 34 6.1 Manuscript I ...... 34 6.2 Manuscript II ...... 35 6.3 Manuscript III ...... 36 7. Conclusions ...... 36 8. Acknowledgments ...... 38 9. References ...... 40

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1. Introduction

1.1 Thesis aim The objectives of this project were: (1) to obtain a precise chronology of distinct breccia samples collected at different Apollo landing sites, (2) perform a thorough petrological and textural investigation of zircon grains, and (3) investigate the possibility of incomplete resetting of the U- Pb system in zircon and phosphates. Each of these three objectives is linked to a more general aim of reassessing the impact history of the Moon, and testing the theory of the “Late Heavy Bombardment” (LHB).

1.2 Early impact cratering (~4.4–3.8 Ga) and the proposed LHB The time interval between the Moon´s formation at approximately 4.5 Ga and the onset of basaltic mare formation at around 3.8 Ga appears to have been characterized by intense bombardment, leading to the formation of the major lunar impact basins (i.e. impact craters with a diameter > 100 km). There is still an ongoing debate whether the impact flux during this period (1) began high and declined steadily during the first ~600 myr, (2) was low but with a cataclysmic increase around 3.9 Ga, (3) increased periodically with several spikes between 4.5–3.9 Ga, or (4) had a moderate increase around 4.1 Ga and declined slowly until 3.9 Ga (Fig. 1).

Prior to the availability of lunar samples, orbital images were used to study the lunar surface. Hartmann (1966) observed that the mare plains are less crater-saturated than the surrounding highlands and concluded that the cratering rate must have been roughly two hundred times higher before the mare formed. Moreover, the impact flux must have decreased rapidly within the first billion years of lunar history. Hartmann (1966, Hartmann et al. 2000) named this period “early intense bombardment” but this term is now outdated. Early analyses of lunar samples lead to a slightly different theory, namely the so-called Late Heavy Bombardment (LHB). Rb-Sr, Ar-Ar, K-Ar, and U-Pb dating of whole rocks from the Apollo 14, 16, and 17 missions yielded a cluster of ages at ~4.00–3.85 Ga (Tera et al. 1974, Turner and Cardogan 1975). These findings were the foundation for the hypothesis that there were several impact spikes with a cataclysmic ending forming the impact basins Imbrium, Orientale, Crisium and possibly other large basins within ~20 myr (Tera et al. 1974, Cohen et al. 2000). This final episode of intense bombardment is referred to as the terminal lunar cataclysm (Tera et al. 1974). Ryder (1990) took this theory even

7 further and argued that the absence of impact melts older than 3.85 Ga implies a low impact flux within the first ~600 myr of lunar history followed by a final intense cataclysm at ~3.85 Ga. Nevertheless, several scientists argued against the theory of a terminal lunar cataclysm and explained the cluster of ages around 3.9 Ga as being the product of sampling bias, since the samples collected during the Apollo and Luna missions were derived from a limited range of areas on the lunar nearside (e.g. et al. 2007). As such, it was suggested that these samples may be dominated by ejecta from only a few impact basins (e.g. Imbrium) and therefore yield radiometric ages that are the same within their analytical uncertainties. Additionally, Hartmann (1975) discussed the possibility that the lunar surface reached crater saturation at ~4.1 Ga and thus, subsequent impacts would have destroyed older impact structures and erased information about pre 4.1 Ga impacts. More recently, Morbidelli et al. (2012) suggested a “sawtooth”-like profile for the lunar impact history. In this model, the impact flux was high but decreased within the first ~400 myr with a moderate uptick at ~4.1 Ga (Morbidelli et al 2012).

Figure 1. Schematic diagram illustrating different models of the early impact flux between the formation of the Moon at approximately 4.5 Ga and 3.7 Ga. The blue line indicates a high cratering rate from the Moon´s formation and a steady decline during the first ~600 myr (Baldwin 1964, Hartmann 1966). Another scenario assumes several impact spikes between 4.5–3.9 Ga with a terminal lunar cataclysm at 3.9 Ga (red dashed line, Tera et al. 1974). The line indicates low impact flux within the first 600 myr and a cataclysmic increase around 3.9 Ga (Ryder 1990). More recently, Morbidelli et al. (2012) proposed a “sawtooth” profile for the lunar bombardment with a moderate increase in impact flux at ~4.1 Ga (yellow line).

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It is believed that the LHB also occurred on Earth and other planetary bodies and therefore, the LHB and a possible terminal cataclysm have important implications for the beginning and evolution of life on Earth (e.g. Zahnle and Sleep 2006). Our current understanding of the early impact flux throughout the early Solar System is however mainly based on investigating lunar samples, because impact structures are much better preserved on the Moon than on Earth. Geological processes, such as or plate tectonics, eliminated old (> 3.0 Ga) impact structures on Earth. In addition, there is a lack of ancient samples from other planetary bodies. Hence, the study of lunar samples offers the wider possibility to investigate the impact flux and the evolution of the early Solar System.

1.2.1 Origin of the LHB Different theories regarding the source of the impactors have been proposed to explain the LHB, but there is no generally accepted explanation. Morbidelli et al. (2001) suggested that a leftover population at the end of planetary accretion caused several spikes in the early impact history, but could not have led to a cataclysm at 3.9 Ga. Another theory states that Uranus and Neptune formed later than the terrestrial planets and Jupiter and Saturn, causing the scatter of icy planetesimals throughout the Solar System (Levison et al. 2001). The late formation of Uranus and Neptune would also have led to the migration of Jupiter and Saturn, resulting in the destabilization of the asteroid belt. Thus, the Moon was bombarded by and planetesimals (Levison et al. 2001). A very similar model was proposed by Gomes et al. (2005), who explain the LHB by the migration of giant planets, which caused the destabilization of the planetesimal disk and the asteroid belt (the so-called Nice model). More recently, the Nice model was taken as a foundation to build the E-belt concept (Bottke et al. 2012). This states that most impactors came from the E-belt, an extended asteroid belt between 1.7 and 2.1 astronomical units from Earth, which is now almost extinct (Bottke et al. 2012). The projectiles were scattered throughout the Solar System due to the late giant planet migration and led to a moderate cataclysm around 4.1 Ga (Morbidelli et al. 2012). The impact rate declined within ~400 myr and a combined bombardment of projectiles from the E-belt and main asteroid belt led to the formation of all lunar basins (Bottke et al. 2012, Morbidelli et al. 2012).

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1.2.2 Recent studies and impact ages Since the early analyses of lunar samples leading to the idea of the LHB and possibly a terminal lunar cataclysm (e.g. Tera et al. 1974), more recent studies have not only utilized new high- resolution remote sensing data, but also updated decay constants for 40Ar/39Ar and Rb-Sr age dating, improved analytical capabilities and, in some cases, new analytical techniques (e.g. Secondary Ion Mass Spectrometry). Additionally, impactite clasts in lunar were analysed in order to gain a better understanding of the early impact history of the Moon. Meteorites provide crucial information since they are derived from random locations of the Moon, possibly also from the lunar farside and they might be composed of source rocks different from those sampled at the Apollo and Luna landing sites. 40Ar/39Ar ages of impactite clasts within lunar meteorites yielded ages around ~4.2 Ga, a peak at ~3.7 Ga, and declining ages to ~2.5 Ga (e.g. Cohen et al. 2005, Joy and Arai 2013). Nevertheless, the younger ages (< 3.7 Ga) cannot be associated with basin forming events since no basins were formed after this time (Joy and Arai 2013). These ages rather reflect localized and/or smaller impact events (Chapman et al. 2007).

The dating of zircon and Ca-phosphates by SIMS analyses of U-Pb isotope systematics has been shown to be of great value when dating impact events (see section 4.1). The U-Pb system of Ca- phosphates is susceptible for thermal resetting due to the low closure around 450– 550 ̊C (e.g. Cherniak et al. 1991). While this trait puts constraints on recording primary crystallisation ages of the analysed phases and their source rocks, it has the potential to date major impact events, which may have reset the U-Pb system in phosphates. For example, the study of Ca-phosphates in lunar breccias collected at the Apollo 14 and 17 landing sites yielded weighted average ages ranging from ~3944 Ma to ~3926 Ma (Nemchin et al. 2009, Merle et al. 2014, Snape et al. 2016a).

While the U-Pb ages of zircon are usually taken as primary crystallisation ages, some recent studies have identified secondary modification through impact events (e.g. recrystallisation or granular zircon). Impact events recorded in lunar zircon were identified at ~4.33 Ga (Grange et al. 2009, Crow et al. 2017), ~4.20 Ga (Pidgeon et al. 2007, Norman and Nemchin 2014, Crow et al. 2017), ~4.11 Ga (Grange et al. 2011), ~3.93 Ga (Grange et al. 2009), and 1.94 Ga (Grange et al. 2013b) with a high degree of confidence. Moreover, in some cases lunar zircon growth within an impact melt could be identified, providing evidence of impact events at ~3.95 Ga, ~3.93 Ga,

10 and ~3.91 Ga (Gnos et al. 2004, Liu et al. 2012, Hopkins and Mojzsis 2015). Even though, these grains represent only a minority among the zircon grains dated, the ages contradict the idea of the most extreme interpretation of the terminal lunar cataclysm (i.e. a low impact flux within the first 600 myr and a cataclysm at ~3.9 Ga; Ryder 1990) or a “sawtooth” profile of the lunar impact flux with a peak at ~4.1 Ga (Morbidelli et al. 2012). The data seem to be in agreement with several spikes in the impact flux throughout the first ~600 myr of lunar history. The ~3.93 Ga ages recorded in several Ca-phosphates and in some zircon grains from several breccias collected at different Apollo landing sites were interpreted to reflect the age of the Imbrium impact (e.g. Liu et al 2012, Snape et al. 2016a). The reoccurrence of this ~3.93 Ga age might indicate that Imbrium ejecta indeed covers large areas of the lunar nearside. Nevertheless, the slightly older U- Pb phosphates ages (3.94 Ga) might originate from a separate major impact event, which would indicate at least two basin forming events within a short time interval of ~10 myr (Snape et al. 2016a). Another possibility is incomplete resetting of the U-Pb system in phosphates, leading to apparently older ages. While the latter possibility cannot be excluded, the additional collection of U-Pb ages of Ca-phosphates and zircon in this study aims to resolve some of the controversies regarding the LHB theory.

2. Geological processes In the following, a short and simplified overview of impact cratering and the magmatic evolution of the Moon are given, as well as an introduction to lunar rocks and minerals. These sections provide information on some basic concepts, background information, and the definition of specific terms, which are used in this thesis. The section about impact cratering aims to explain different types of impact craters and the distinct stages during an impact event. The following introduction to the magmatic evolution is important in order to understand the diversity of lunar rocks and the chemical and compositional variability of lunar breccias. Section 2.3 introduces lunar minerals and rocks, some of which are found in the breccias analysed in this study.

2.1 Impact cratering Lunar impact craters are the result of a meteoroid, asteroid, or striking the lunar surface (i.e. an impact event) and leaving a bowl-shaped depression. The impact event itself can generally be divided into three stages: the contact and compression stage, the excavation stage, and the modification stage (Gault et al. 1968, Fig. 2). When the impactor hits the target (e.g. the

11 lunar surface), shock waves are generated, which travel into the target and backwards into the projectile. The shock wave causes particle motion and most of the initial kinetic energy of the projectile is transferred to the target, leading to compression of the underlying rocks and acceleration of material (Melosh 1989). The compressions stage is characterized by extremely high pressures (several hundred GPa) and (> 10000 K), which results in the melting and vaporization of large volumes of the target and the entire projectile (Melosh 2012). The excavation stage starts directly after the compression stage and comprises the propagation of the shock waves into the target, where they expand and their intensity decreases. As the shock waves reach free surfaces (e.g. the ground surface), rarefaction wave are released, leading to the formation of a flow field (i.e. particle motion) and initially the crater excavation. The excavation ends with the formation of a transient crater, which is modified during the modification stage. The modification stage is characterized by material inwards into the crater and filling the crater floor. The crater walls of larger craters collapse forming terraces and central peaks rise in the centre (Melosh 1989). The final crater is typically about ten times larger than the impactor.

The material thrown out during the excavation stage is called ejecta and consists of impact breccias, i.e. the rocks formed during an impact event. The ejecta usually forms continuous and discontinuous deposits as well as ejecta rays. The continuous ejecta usually builds an uninterrupted circular around the and it extends approximately one crater radius from the crater rim (Melosh 1989). In contrast, the discontinuous ejecta builds patchy and localized deposits further away from the impact site. During the ejecta deposition larger blocks from deeper target material are deposited closer to the crater rim, while the finer grained material is deposited further away from the impact crater (Melosh 1989). Ejecta rays may occur beyond the discontinuous ejecta deposit as elongated streaks. Secondary craters can occur behind the continuous ejecta blanket and often form clusters or crater chains. The process of secondary cratering leads to extensive mixing of local material into the ejecta blanket. Ejecta deposits from large impact basins may therefore be completely dominated by local material (Oberbeck et al. 1975).

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Figure 2. Schematic drawing of the different stages during an impact event, leading to the formation of a (see text for more details; modified after B.M. French and D. A. Kring/LPI/UA).

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Based on their morphology, impact structures can be divided into three categories: simple craters, complex craters, and basins. Simple craters are usually bowl-shaped, with a flat or rounded floor and smooth rims (Melosh 2011). If the diameter of simple craters increases, they evolve into more complex craters with terraced rims and a flat floor with an uplifted central peak. This transition takes place in the 15–20 km diameter range for craters on the Moon (Pike 1977). Once the diameter exceeds 100 km, peak rings appear on the crater floor and mark the transition to impact basins (Hartmann and Kuiper 1962), which can be further divided into central peak basins (with peak rings surrounding a central peak), peak-ring basins (which lack a central peak) and multiring basins (which have up to six concentric rings). The latter are usually more than 400 km in diameter. The diameters mentioned above refer to the final crater diameters and not to the transient crater diameters.

2.2 Magmatic evolution and major lunar provinces It is believed that the Moon was in a nearly completely molten state after its accretion and that it was covered by a so-called (LMO; Wood et al. 1970). This concept was developed after the first sample returns from the Apollo missions, which indicated that the Moon has a plagioclase-rich crust. Since plagioclase is a relatively light mineral, it was assumed that it floated on top of a magma ocean (Wood et al. 1970). In this model, as the LMO began to cool, olivine and pyroxenes, as well as ilmenite crystallised and due to their higher density, sank to the bottom of the LMO. This was followed by (plagioclase feldspar) crystallisation. Anorthosite has a lower density and therefore, floated to the surface, where it produced a plagioclase-rich crust (Warren 1985). After ~99% of the magma ocean crystallised, a residual melt enriched in trace elements was left over and termed urKREEP (where KREEP is an acronym referring to the presence of Potassium, Rare Earth Elements, Phosphorus; Warren and Wasson 1979). The urKREEP layer was thought to be located between the upper anorthositic crust and the mafic mantle and to have initially an evenly global distribution with a thickness of approximately 2 km (Warren and Wasson 1979). Several questions regarding the LMO theory still remain, e.g. the timing and duration of the LMO, its depth, and whether the crystallisation was fractional or in equilibrium.

KREEP-rich or KREEPy material is found in many breccia samples and in some from the and 17 landing sites (e.g. Ryder et al. 1977, Ryder 1987). These basalts are thought to

14 have formed by partial melting of the upper mantle mixed with urKREEP (Wieczorek and Phillips 2000). Nevertheless, urKREEP has not been sampled yet by any mission and KREEP basalts show unique chemical compositions different to the supposed composition of the urKREEP layer. Based on more recently obtained remote sensing data from the Clementine and Lunar missions, scientists discovered an asymmetrical distribution of KREEPy material, with localised concentrations around the Imbrium basin and the Oceanus Procellarum region. This region was therefore termed the Procellarum KREEP terrane (PKT; Jolliff et al. 2000). Furthermore, Jolliff et al. (2000) divided the lunar crust in at least three major provinces based on geochemical, geophysical, and petrologic data: the PKT, the Feldspathic Highlands Terrane (FHT), and the South Pole- Terrane (SPAT). The FHT is characterised by an elevated topography, a highly feldspathic composition, and it is heavily cratered. The FHT covers about 60% of the lunar surface. The SPAT is defined by the perimeter of the South Pole-Aitken Basin. This ~2600 km diameter feature is the largest impact basin in the Solar System. Moreover, it exhibits relatively high FeO concentrations compared to non-mare crust and a lack of mare volcanism compared to the nearside. The latter would have been expected due to its depth (up to 12 km below the surrounding highlands) and thinning of the crust (Jolliff et al. 2000).

2.3 Lunar minerals and rocks The study of lunar minerals is crucial in order to understand the chemical and physical conditions under which lunar rocks formed (e.g. the temperature or the pressure). Lunar minerals formed in a low- and highly reducing environment, which is why metallic iron (Fe) and (FeS) are common in lunar rocks (Papike et al. 1991). One of the big differences between terrestrial and lunar rocks is the lack of water-bearing minerals in lunar rocks, such as , oxyhydroxides, or amphiboles. These minerals may occur near the poles, where H deposits were detected via remote sensing (e.g. Feldman et al. 2001, Li et al. 2018). Also silica minerals (SiO2) and potassium feldspar (KAlSi3O8), which are abundant on Earth, are rare in lunar samples (Papike et al. 1991). 2- 3+ Minerals containing carbonate (CO3 ) and ferric iron (Fe ) are lacking on the Moon. The most abundant minerals in lunar rocks are silicate minerals, such as pyroxene ((Ca, Fe Mg)2Si2O6), plagioclase feldspar (NaAlSi3O8–CaAl2Si2O8), and olivine ((Mg,Fe)2SiO4). Furthermore, oxide minerals are common in lunar rocks, especially in mare basalts. Examples of these include ilmenite ((Fe, Mg)TiO3) or armalcolite ((Fe, Mg)Ti2O5), the latter having been first described in

15 lunar rocks (e.g. et al. 1970) and named after the astronauts , , and .

Lunar rocks can be divided into different groups: (1) pristine highland rocks, (2) pristine basaltic volcanic rocks, (3) impact breccias (discussed in chapter 4.1), and (4) (Taylor et al. 1991). Pristine rocks are by definition rocks that consist of only a single rock type (i.e. no evidence for petrographic or chemical mixture of different rocks and absence of siderophile elements). Nevertheless, these ancient pristine non-basaltic rocks are rare among the lunar samples due to the melting, modification, and mixing of rocks through intense meteoroid bombardment of the lunar surface (Taylor et al. 1991). Therefore, the majority of samples derived from the lunar highlands are polymict impact breccias. Pristine highland rocks can further be divided into ferroan , Mg-suite rocks, and KREEP rocks. The ferroan anorthosites crystallised as coarse-grained igneous rocks and consist mainly of plagioclase (> 90%), as well as minor amounts of pyroxene and olivine. Mg-suite rocks comprise a range of distinct rocks and the compositional range is not accurately defined, but they are usually more mafic in composition than ferroan anorthosite rocks (Taylor et al. 1991). Rocks included in this group are, among others, norites, troctolites, and ultramafic rocks (i.e. consisting mainly of olivine and/or pyroxene). Almost all KREEP rocks are basalts (from the Apollo 15 landing site), but a few felsites and quartz monzodiorites with KREEPy composition have been found as clasts within breccias collected at the Apollo 14 and 15 sites (e.g. Warren et al. 1983, Jolliff et al. 1991).

Basaltic volcanic rocks comprise basaltic lavas/flows and pyroclastic deposits. Basaltic mare basalts cover ~17% of the lunar surface (e.g. Head and Wilson 1992), but are highly concentrated on the lunar nearside, where they fill lunar impact basins (and form so-called maria). Mare basalts are thought to have originated from partial melting of the mantle and can be divided into different groups based on their Ti content - High-Ti, Low-Ti and Very-low Ti (Papike et al. 1976, Papike and Vaniman 1978).

The entire lunar surface is covered by a layer of unconsolidated material, the lunar regolith. This regolith layer is the result of continuous impact cratering and consists of breccia fragments, mineral fragments, pristine rock fragments, as well as and agglutinates (McKay et al. 1991). The particles are usually < 1 cm in size. In mare regions, the regolith is thought to have a

16 thickness of ~2–4 m, and may reach a thickness of 6–8 m on the farside and non-mare nearside regions (Bart et al. 2011).

3. Geological setting The samples used in this study were collected during the Apollo missions in the late 1960s and the early 1970s. While the Apollo astronauts underwent intense geological training before their missions, in most cases it is difficult to put the collected samples into their geological context. This is primarily due to the fact that the lunar surface is covered by a regolith layer from which the samples were taken and not from actual bedrock exposures. Hence, the context remains poorly understood in most cases and the link between the U-Pb ages obtained in this study and specific impact events remains speculative. Nevertheless, high spatial resolution remote sensing data, the chemistry, and detailed petrological investigations of the samples, as well as models to calculate the ejecta distribution of large impact events can help to determine likely origins of the analysed samples or at least restrict them to a few known impact events. In this thesis, Ca- phosphates and zircon in thin sections from breccias collected during the Apollo 12, 14, and 17 missions were investigated (Fig. 3).

Figure 3. Lunar Reconnaissance Orbiter Camera (LROC) image of the lunar nearside (NASA/GSFC/Arizona State University). Breccias analysed in this study were collected at Apollo 12, 14, and 17 landing sites (yellow stars). Highlighted are also some of the major lunar impact basins.

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3.1 Apollo 12 The Apollo 12 landing site was chosen to demonstrate pinpoint landing capability, which was successfully accomplished when the lunar module landed on the 19th November 1969 within 200 m of the Surveyor 3 landing site. The landing site is located in the southeastern part of Oceanus Procellarum (3.2S, 23.4W) on the northwestern rim of the Surveyor crater (Fig. 4). It is also situated on a bright ray of the Copernicus crater, a relatively young crater (~800 Ma, Wilhelms 1987) approximately 370 km north of the landing site. Two extravehicular activities (EVAs) were performed during which 34.3 kg of samples were collected, mostly basalts and a few impact breccias. The landing site is covered by Eratosthenian mare basalts, which can be divided into four distinct types: olivine basalts, pigeonite basalts, ilmenite basalts, and a single fragment of feldspathic basalt (Neal et al. 1994). Radiometric ages for these basalts range from 3129 ± 10 Ma to 3242 ± 13 Ma (Snape et al. 2016b). These basalts are overlain by a thin regolith layer (< 3 m), which contains non-mare KREEP-rich material, probably incorporated into the regolith by lateral transport of different impact events and vertical mixing (Jolliff et al. 2000, Korotev et al. 2011). Breccia 12013, sampled during the mission, was analysed in this study; however, the exact sampling location is unknown.

Figure 4. Traverse map of the Apollo 12 landing site. Sample collection locations are displayed by white circles. Also shown are the location of the Surveyor 3 lander (yellow circle), the Lunar Module (LM; blue circle), and the Apollo Lunar Surface Experiment Package (ALSEP; blue circle). Black lines indicate the extravehicular activity (EVA) periods. Background image is from the LRO Narrow Angle Camera, Frame M175428601R (NASA/GSFC/Arizona State University).

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3.2 Apollo 14 The Apollo 14 landing site (3.7S, 17.5W) is located approximately 1100 km away from the centre of the Imbrium basin and 550 km south of its outer rim, just 1110 m west of the Cone Crater (Fig. 5). The site was selected to sample the formation, which was assumed to consist of material excavated by the Imbrium impact (e.g. Swann et al. 1971). Therefore, the landing site offered the possibility to date one of the major lunar impact events and an important calibration point for the lunar cratering chronology. However, studies of the samples revealed that the breccias most likely represent a complex mixture of local pre-Imbrian material and Imbrian ejecta (Morrison and Oberbeck 1975, Hawke and Head 1977, Stöffler 1989). The 42.3 kg of rocks collected during the mission include impact-melt breccias, clast-poor impact melts, and fragmental breccias and are in general KREEP-rich. Cone Crater is one of the major geologic features of the landing site, approximately 340 m in diameter and ~75 m deep (Wilhelms et al. 1987). It is believed the impact might not only have excavated material from the Fra Mauro Formation, but also blocks of bedrock material from below the formation (Head and Hawke 1975, Stöffler 1989, Snape et al. 2016a). However, this depends on the thickness of the formation on the landing site, which was initially estimated to be 100–200 m (Eggleton and Offield 1970), but subsequent seismic studies showed that the thickness is more likely to be between 16–76 m (Watkins and Kovach 1972). Four samples were analysed in this study (breccias 14305, 14306, 14314, and 14321).

Figure 5. Map of the Apollo 14 landing site. Indicated are the sample locations from which the samples investigated in this study were collected (red circles; white circles are other sample locations). Also displayed are the EVAs (black lines), the location of the LM (blue circle) and some craters in the vicinity of the landing site (white dotted lines). The background image is from the LRO Narrow Angle Camera, Frame M102265088L (NASA/GSFC/Arizona State University).

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3.3 Apollo 17 The Apollo 17 landing site was located in the Taurus-Littrow valley, which is close to the southeastern rim of the Serenitatis impact basin (20.2N, 30.8E). It was chosen since it was assumed that the astronauts could sample older (pre-Imbrian) material from the highland massifs and possibly young volcanic material. The area around the landing site can be divided into different geological units like the high massifs (North and South massif), the so-called Sculptured Hills, and mare basalts which cover the valley floor (Fig. 6a). The Sculptured Hills are referred to as a “knobby” unit found near the massifs of the Apollo 17 landing site and were initially thought to represent impact ejecta from the Serenitatis impact event (Head 1974). Later studies revealed however, that the Sculptured Hills overlie large crater rims, which are superposed on the Serenitatis basin rim and are therefore younger than the Serenitatis impact event (Spudis and Ryder 1981). Spudis and Ryder (1981) also noted that lineations in the Sculptured Hills Terrane points towards the Imbrium basin. Moreover, the knobby hills appear not only at and around the Apollo 17 site but are widespread throughout the Taurus Mountains and occur in the highlands between the Serenitatis and Crisium basins (Spudis et al. 2011). These discoveries led to the idea that the Sculptured Hills might have originated from the Imbrium impact event and not from the Serenitatis impact. The high massifs at the landing site are thought to have been lifted up by the Serenitatis impact event (Head 1974) and reach heights up to 2300 m. Their composition is a mixture of feldspathic granulitic material and noritic impact melt in addition to high-Ti basalts on flanks of the massifs at low elevations and pyroclastic deposits at high elevations (Robinson and Jolliff 2002). In total, 110.5 kg of samples were collected during this mission. Four breccia samples (72255, 76015, 76055, and 76215) were investigated in this study and they were collected from boulders at the North and South massif (Fig 6a). These boulders can be traced back to their sources using new remote sensing data from the Lunar Reconnaissance Orbiter Camera (LROC; Hurwitz and Kring 2016; Fig. 6b&c).

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Figure 6. (a) Traverse map of the Apollo 17 landing site indicating sample locations from breccias analysed in this study (red dots). Displayed are also the lunar rover vehicle (LRV) sampling and traverse stations (black dots), the EVA tracks (white lines), the landing site (blue circle), and some impact craters at the landing site (white circles). The background image is from LRO NAC, frame M104318871 (NASA/GSFC/Arizona State University). (b) NAC image (Frame M134991788R; NASA/GSFC/Arizona State University) of a boulder that rolled down the slopes at the North Massif and left a trail. (c) Astronaut Jack Schmitt next to the large boulder at Station 6. The image is Apollo image AS17-140-21496.

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4. Methods

4.1 Impact Breccias In this study, thin sections of impact breccias sampled at different Apollo landing sites were investigated. Impact Breccias are rocks that are formed during meteoroid or asteroid impacts and are made up of different type of rock fragments, , and mineral clasts, which were excavated and mixed during the impact. These clasts and fragments are surrounded by a so-called matrix. The matrix is composed of a fine-grained material, which is either a solidified melt (crystalline or glassy), clastic (fragmental), or a mixture of both (Taylor et al. 1991). The larger rock clasts may have been derived from a single or different target rocks, this distinction leads to a subdivision of breccia samples between those that are referred to as monomict and those referred to as polymict. Monomict breccias are composed of a single target rock, which has been crushed into numerous smaller fragments. On the other hand, polymict breccias consist of different types of material (e.g. different types of bedrock, older impact breccias, impact melts), which were mixed and potentially melted down during the impact. These polymict breccias can be further divided into distinct breccia types: fragmental breccias, glassy melt breccias, crystalline melt breccias, clast- poor impact melts, granulitic breccias, dimict breccias, and regolith breccias (Taylor et al. 1991). Moreover, the breccias then can be classified based on their texture, like e.g. poikilitic (small crystals of one mineral enclosed in another mineral), subophitic (pyroxene partially encloses plagioclase) or aphanitic (fine-grained). These textures depend on the environment where the minerals formed; e.g. an aphanitic texture indicates rapid cooling. Clast-poor impact melt breccias are difficult to distinguish from pristine igneous lunar rocks (derived from the lunar interior), since their texture is essentially the same. The best way to distinguish between them is to analyse their siderophile (“iron-loving”) element concentrations. When compared to pristine rocks, impact melt breccias are highly enriched in siderophile elements due to meteoroid contamination (e.g. Warren and Wasson 1977). The highly siderophile element (e.g. Iridium, Gold, Osmium) concentrations can also be used to identify the projectile, such as ordinary or iron meteorites (e.g. Tagle and Hecht 2006).

4.1.1 Ca-Phosphates

Merrillite (“lunar whitlockite”, Ca18Na2Mg2(PO4)14) and apatite (Ca5(PO4)3(F,Cl,OH) are common accessory minerals in lunar rocks and meteorites (Papike et al. 1991). Compared to terrestrial whitlockite (Ca18Mg2(PO4)12[PO3(OH)]2), lacks hydrogen in its crystal

22 lattice. Lunar apatite contains very little OH and is instead generally Cl- and F-rich. Apatite and merrillite are in general found in KREEP-rich rocks (Smith and Steelen 1976) and lunar merrillite has a high Y+REE concentration compared to Martian merrillite (Jolliff et al. 1993, 2006). Phosphates have a low U-Pb closure temperature (i.e., when there is no more diffusion of the parent or daughter isotope into and out of the crystal lattice) around 450–550˚ C (Cherniak et al. 1991), which makes their U-Pb system susceptible to complete resetting during a thermal event. This attribute makes phosphates important candidates for recording thermal and/or impact events. The complete resetting of the U-Pb system in phosphates during a thermal event is a key assumption in this study, because the U-Pb ages of Ca-phosphates were used to obtain impact ages of distinct breccias from different Apollo landing sites. Since their ages should represent the age of the thermal resetting event, it makes no difference whether the phosphates crystallised in the impact melt or whether they occur within lithic clasts (and thus crystallised in different source rocks pre-breccia formation). Ca-phosphates analysed in this study occur sometimes as composite apatite-merrillite grains (Fig. 7a) and their texture ranges from euhedral (well-formed crystal faces, i.e. hexagonal in the case of apatite, Fig. 7b), subhedral (partly bound by crystal faces), to anhedral (no well-formed crystal faces).

Figure 7. Backscattered electron (BSE) images of two Ca-phosphate grains analysed in this study. (a) A poikilitic and irregular shaped composite merrillite-apatite grain in breccia 12013 (Apollo 12). (b) A euhedral apatite grain located in a void in breccia 76215 (Apollo 17). The dark areas are the rastered areas and Secondary Ion Mass Spectrometry (SIMS) spots.

4.1.2 Zircon (ZrSiO4) Zircon is a late-stage accessory mineral, which forms in melts saturated in Zr. The residuals of the proposed LMO were enriched in incompatible elements, such as K, Rare Earth Elements, and K (KREEP), as well as Zr. These magmas are thus thought to be the source for lunar zircon (Dickinson and Hess 1982). Zircon incorporates Th and U into its crystal lattice, while Pb2+ is

23 highly incompatible and is therefore only incorporated in very small amounts. This attribute, together with its high closure temperature of the U-Pb system around 950–1100 ˚C (e.g. Cherniak and Watson 2000), and its resistance to thermal disturbance and weathering makes zircon one of the most important minerals for radiometric dating. Hence, the U-Pb study of lunar zircon has been shown to be particularly useful in determining crystallisation ages of rocks. On the downside, zircon grains occur often as small (< 20 µm) isolated fragments within the breccia matrix, making it difficult to establish a link to their original source rocks. This missing link puts constraints on the interpretation of their U-Pb ages, since they might reflect magmatic crystallisation, impact, or partially reset ages. Nevertheless, zircon can exhibit primary and secondary microstructures, which can provide important information when interpreting the obtained U-Pb ages (primary crystallisation ages vs. secondary ages or partially reset ages, e.g. Grange et al. 2013a). Primary microstructures, such as oscillatory or sector zoning seen in Cathodoluminescence, are usually taken as indicator for magmatic crystallisation. Some grains occurring within lithic clasts might show petrological evidence that they crystallised within these clasts (e.g. poikilitic zircon growth, Fig. 8a). Moreover, if several grains occur within the same lithic clast and yield an intra-grain and inter-grain age consistency within the same lithology, these ages can be taken as primary crystallisation ages and thus, as age of the clast. In a few cases, zircon growth from an impact melt could be identified based on their distinctive texture (i.e. poikilitic zircon grains found in e.g. Apollo 12 breccias, Liu et al. 2012 or Apollo 14 breccias, Hopkins and Mojzsis 2015).

On the other hand, secondary microstructures like recrystallisation, planar deformation features (PDFs), reaction structures, granular zircon (Fig. 8b), or crystal-plastic deformation indicate a secondary event, which modified the zircon grains. These secondary events can lead to the formation of new zircon, and partially or completely reset the U-Pb system in the primary zircon. PDFs, for example, might enable fast Pb diffusion through the grain, leading to partial Pb loss (Grange et al. 2013a). If the U-Pb system was partially reset, the grains often do not show any inter-grain age consistency or parts of the grain yield younger U-Pb ages. As a result, only the oldest ages of these grains can be taken as minimum crystallisation age and the younger ages are interpreted as partial reset ages. In some cases, it is possible to distinguish between the primary crystallisation age and the secondary event. An example is the “tiger” grain (Apollo 17 impact melt breccia 73235,59; Grange et al. 2011), where the primary zircon grain yielded a 207Pb/206Pb

24 age of 4354 ± 8 Ma and a recrystallised part of the grain a younger 207Pb/206Pb age of 4106 ± 18 Ma (Grange et al. 2011).

Figure 8. BSE images of two zircon grains. (a) A poikilitic zircon grain with straight to rounded and embayed crystal boundaries. This grain is located within a lithic clast and interpreted to have crystallised in the clast. (b) A baddeleyite grain (Bdl) surrounded by granular zircon (Zrn) located in the brecciated matrix of sample 14314 (Apollo 14). 4.2 U- Pb geochronology The U-Th-Pb decay system is unique due to the occurrence of three independent decay processes. Lead has four naturally occurring stable isotopes: 204Pb, 206Pb, 207Pb, and 208Pb. The parent isotopes 235U, 238U, and 232Th decay, via a series of unstable intermediate daughter isotopes, to their respective stable daughter isotopes 207Pb, 206Pb, and 208Pb. 204Pb is the only non-radiogenic isotope and can, therefore, be used to correct for Pb contamination from other sources. The half- lives of the intermediate daughter isotopes are much shorter than that of the parent isotopes, which have half-lives of 4.47 b.y. (238U), 0.704 b.y. (235U), and 14.01 b.y. (232Th; Jaffey et al. 1971). The decay chains for the two uranium isotopes (which are most relevant for this study) can be expressed by the following age equations:

206 206 238 ʎ238t PbP = PbI + U (e – 1) [1]

207 207 235 ʎ235t PbP = PbI + U (e – 1) [2]

, where the subscript P indicates the abundance of the isotope at present time, and I indicates the initial abundance of that isotope (also referred to as common Pb), t is the time since the system 238 235 was closed, and λ238 and λ235 are the decay constants of U and U (e.g. Steiger and Jäger 1977). Isotope ratios can be measured much more precisely compared to their concentrations, which can be achieved by dividing equations [1] and [2] by 204Pb:

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206푃푏 206Pb 238푈 ( ) = ( ) + ( ) (푒휆238푡 − 1) [3] 204푃푏 204푃푏 204푃푏 퐼

207푃푏 207푃푏 235푈 ( ) = ( ) + ( ) (푒휆235푡 − 1) [4] 204푃푏 204푃푏 204푃푏 퐼

If the contribution of initial lead is marginal or has been corrected for, the equations [3] and [4] can be simplified to:

206푃푏∗ ( ) = (푒휆238푡 − 1) [5] 238푈

207푃푏∗ ( ) = (푒휆235푡 − 1) [6] 235푈

, where the asterisk stands for radiogenic lead accumulated through the decay of their parent isotopes. The dual decay system allows a third isochron equation by dividing equation [4] by equation [3]:

207푃푏 207푃푏 ( ) − ( ) ∗ 204푃푏 204푃푏 235푈 (푒휆235푡 − 1) 207푃푏 퐼 = ( ) = ( ) [7] 206푃푏 206푃푏 238푈 (푒휆238푡 − 1) 206푃푏 ( ) − ( ) 204푃푏 204푃푏 퐼

This equation permits an age determination without the need to measure the U concentration, because the 235U/238U is assumed to be a known constant of 137.88 (e.g. Steiger and Jäger 1977) in meteoritic and terrestrial systems. More recently, a value of 137.818 ± 0.045 (2σ) was suggested for zircon geochronology (Hiess et al. 2012). Nevertheless, this new value results in a shift of age of only ~0.2‰ (~0.7 Ma) at 3.9 Ga (Hiess et al. 2012), which is within the analytical uncertainties. The value of 137.88 was used in this study (Isoplot version 4.15, Ludwig 2012). U- Pb data can be displayed in different ways, but the most common ways are the conventional (Wetherill 1956) and the inverse concordia (Tera and Wasserburg 1972) diagrams. These diagrams help to visualize recent and ancient Pb loss.

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4.2.1 Wetherill concordia diagram In the Wetherill concordia diagram, the 206Pb/238U ratio is plotted versus the 207Pb/235U ratio (Fig. 9). The concordia curve is the locus of all points with equal 206Pb/238U and 207Pb/235U ages. All points falling on the concordia curve are called concordant and indicate a closed system, whereas points lying off the curve are called discordant and suggest open system behaviour. Discordance can be caused by Pb and/or U loss or gain or by analysing a mixture of older and younger domains.

Figure 9. Schematic Wetherill concordia diagram. Data points falling on the concordia are termed concordant and data points off the concordia are discordant (black dots). In this example, the upper intercept at 4.0 Ga (blue dot) is the true formation age, whereas the lower intercept at 1.5 Ga (yellow dot) corresponds to the time of the Pb-loss event.

4.2.2 Tera-Wasserburg diagram The Tera-Wasserburg diagram plots the 207Pb/206Pb ratio against the 238U/206Pb ratio and was invented by Tera and Wasserburg (1972), since they noticed a disagreement of U-Pb ages and Rb-Sr or K-Ar ages of lunar rocks. The discrepancy is caused by the incorporation of excess radiogenic Pb at the time of crystallisation (Tatsumoto 1970). Just like the Wetherhill concordia diagram, all points on the curve are concordant and the ones falling off the curve indicate disturbance of the U-Pb system. In a simple case, all analyses will form a discordia line and the intersections with the concordia curve will define the time of crystallisation and disturbance,

27 respectively. One advantage of the Tera-Wasserburg diagram is that recent Pb loss is easily to detect since it will result in a horizontal displacement of points from the concordia curve (Fig. 10).

Figure 10. Schematic Tera-Wasserburg diagram, where the 207Pb/206Pb ratio is plotted against the 238U/206Pb ratio. Also in this case, the blue dot indicates the crystallisation age and the yellow dot the Pb loss event. 4.3 Initial lunar Pb The 207Pb/206Pb ages have to be corrected for common terrestrial Pb and initial lunar Pb, since minor amounts of either of them could lead to artificially older 207Pb/206Pb ages. Common terrestrial Pb is thought to be derived from the sample preparation and polishing (Nemchin et al. 2009), whereas initial lunar Pb was incorporated into the minerals during crystallisation. Compared to Earth, lunar rocks generally have relatively low Pb concentrations. It is assumed, that the Moon lost most of its Pb during its formation (around ~4.5 G) as a result of a giant impact. The loss of Pb resulted in a high 238U/204Pb ratio (µ) of many lunar rocks (~100–600) compared to the µ-values obtained for the Earth´s mantle (~8–10; Kramers and Tolstikhin 1997). Therefore, lunar rocks have a highly radiogenic Pb composition. Moreover, as mentioned previously, phosphates and zircon incorporate little or no Pb during crystallisation and thus, the initial Pb contribution is assumed to be minor and has little effect on the ages determined.

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This assumption can be tested, by plotting the uncorrected 207Pb/206Pb ratios against the 204Pb/206Pb ratios, which is shown in Figure 11. In this example, the uncorrected 207Pb/206Pb ratios and 204Pb/206Pb ratios of phosphates from one Apollo 17 breccia were taken. These grains exhibit highly radiogenic compositions with very low 204Pb/206Pb ratios, and the trend of the analysed grains points towards modern terrestrial Pb (Stacey and Kramers 1975), indicating that the samples were likely contaminated by common terrestrial Pb during sample preparation. Thus, the model of Stacey and Kramers (1975) and their values for present-day terrestrial Pb isotopic ratios were used to correct the data for the effects of common Pb contamination.

Figure 11. An example of uncorrected 207Pb/206Pb and 204Pb/206Pb ratios of analysed Ca-phosphates from one Apollo 17 breccia. The circle indicates the primordial Pb isotope composition of the Solar System as derived by the Canyon Diablo Troilite (CDT) iron . The triangle indicates the present day terrestrial common Pb (S-K terrestrial Pb, Stacey and Kramer 1975). The three growth curves (dashed lines) display different 238U/204Pb (µ) values, corresponding to the composition of initial lunar Pb originating from CTD. The growth curves assume a formation age of the Moon at ~4.5 Ga and a single stage model of lunar Pb isotopic evolution. The black dots (labelled t1 and t2) mark the intercept between the growth curves and the isochrones and indicate the composition of initial lunar Pb at two randomly chosen times. 4.4 Partial resetting One of the key assumptions in this study is the complete resetting of the U-Pb system in phosphates during an impact event, which means that the in situ accumulated Pb from the decay of U is completely lost and thus, the radiometric clock is reset to zero. However, there is a possibility that some phosphates and zircon grains experienced partial Pb loss, leading to

29 artificially, and therefore meaningless, older 207Pb/206Pb ages. While Pb loss is relatively easy to detect in younger grains (< 4.0 Ga) due to visible discordance of the data, it becomes more difficult in case of old (> 4.0 Ga) grains (and the majority of lunar zircon and phosphate grains are older than 3.9 Ga). The difficulties arise from the fact that any ancient disturbance (> 4.0 Ga) would result in a scatter of 207Pb/206Pb ages, but the data points would remain concordant within the analytical error due to the uncertainties on the U/Pb ratios at this time. Partial resetting can be detected if multiple analyses of a single grain result in a range of 207Pb/206Pb ages, which cannot be combined to give a weighted average age. Only the oldest age of these grains can be taken as minimum crystallisation age. However, even the minimum age might be millions of years younger than the actual crystallisation age. Equally, it is possible that the youngest age might just give a maximum limit for the secondary event. Zircon grains can exhibit specific microstructures, indicating a secondary event which may have led to partial resetting (e.g. recrystallisation areas, neoblasts, amorphous domains; e.g. Grange et al. 2013a). As an example, the “pomegranate” grain (Smith et al. 1986, Pidgeon et al. 2007) consists of a primary zircon grain, which has been broken into numerous fragments and shows sector zoning in CL. The primary zircon is surrounded by secondary zircon, which is dark in CL and has high U and Th contents. In this case, the secondary zircon gave an age of 4187 ± 11 Ma, which is 130 myr older than the primary zircon (Pidgeon et al. 2007).

5. Analytical methods

5.1 Scanning Electron microscopy (SEM) A scanning electron microscope can be used to obtain information about the chemical composition, texture, and crystalline structure of samples. A focused beam of is scanned over the sample, causing the emission of back scattered and secondary electrons, x-rays, and visible light (Cathodoluminescence). By rastering the electron beam, the data obtained are combined into images. Secondary electron images show the morphology and topography of the samples, while backscattered electron images (BSE) contain contrast information of areas with distinct chemical compositions. Heavier elements with a high atomic number will appear brighter than elements with a low atomic number. X-rays are used for the elemental analysis and diffracted backscattered electrons help to investigate crystal structures and the orientation of minerals. In this study, the SEM analyses were carried out with a Quanta 650 FEG-SEM and

30 accompanying Oxford Instruments Energy Dispersive Spectroscopy (EDS) detector at the Swedish Museum of Natural History.

5.1.2 Cathodoluminescence (CL) imaging A Cathodoluminescence (CL) detector can be attached to the SEM to produce CL images. These images provide important information about the internal crystal structure and its growth history. For example, CL images of zircon grains can exhibit characteristic oscillatory or sector zoning, or reveal recrystallisation areas (Fig. 12). These different domains might yield distinct ages and therefore, primary zircon growth can be distinguished from secondary modification. Varying abundances of uranium (U) and yttrium (Y) within the crystal cause the apparent zoning in zircon (Boggs and Krinsley 2006). There is a negative correlation between the U and Y contents and the Cl emission (Rubatto and Gebauer 2000). If a CL zone appears darker, is enriched in U and Y (Hanchar and Rudnick 1995, Rubatto and Gebauer 2000). Moreover, crystal-plastic deformation and radiation damage cause a decrease in CL emission (Timms and Reddy 2009). The CL images of zircon grains in this study were obtained using the ChromaCl2 system.

Figure 12. Cathodoluminescence (CL) images of two zircon grains. (a) This zircon grain exhibits oscillatory zoning in CL. (b) Zircon grain with a brighter CL zone at the grain boundary. 5.2 Secondary ion mass spectrometry (SIMS) SIMS is a powerful and one of the most important techniques for in situ analyses of solid materials. It offers the possibility to conduct measurements at a high-spatial resolution and therefore, can potentially allow for multiple analyses on a single mineral grain. The main components of a SIMS are the primary ion source, the primary column, secondary ion extraction and transfer, the , and the mass and secondary ion detectors. A beam of primary is generated by an ion source and accelerated onto the sample, causing the ejection of

31 secondary ions from the sample surface, a process known as sputtering (Fig. 13). This primary + - - ion source creates an ion beam of either cations (e.g. Cs ) or anions (e.g. O or O2 ) to eject ions that are generally of opposite charge. The most common primary beam species are Cs+ and O- since they are highly electropositive and electronegative, respectively (Ireland 1995). A negative - + + ion beam (O2 ) was used in this study to emit the secondary ion species (Pb and UO ). In order to avoid charging of the sample surface, the samples are coated with a thin conductive layer (e.g. a 30nm gold layer was used in this study). The majority of ejected particles are neutrals and only a minor amount is ionized (Ireland 1995). The secondary ions are accelerated into the mass spectrometer, where they are separated by a magnetic field according to their mass-to-charge ratio (Reed 1989). If another ion has the same nominal mass as the ion of interest, mass interferences can occur. This interference can be resolved with a sufficient mass resolution; e.g. M/∆M > 5000 for Pb in zircon in order to separate Pb peaks from isobaric interferences (Ireland and Williams 2003). The number of secondary ions can then be counted by different ion detectors (e.g. an analog charge measuring cup for large signals or a very low noise, ion counting electron multiplier) and compared to the count rates of a known reference material. These reference materials should have the same (or similar) crystal structure and are used to correct for so-called matrix effects, i.e. differences in the relative emission of elemental and/or isotopic ionic species resulting from differences in the chemistry and structure of the sample. In this study, Ca- phosphate analyses were corrected against the 2058 Ma apatite standard BR2 (Grange et al. 2009) and zircon analyses were calibrated against the 1065 Ma zircon crystal 91500 (e.g. Wiedenbeck et al. 1995). Due to the lack of a merrillite standard, the 2058 Ma apatite standard BR2 was also used as a reference for merrillite analyses. Minor differences in mineral chemistry therefore can, in this case, lead to minor uncertainties. Two analytical modes are available to collect the ions: either monocollector (peak-hoping mode, where different masses are detected sequentially) or multicollector (multiple isotopes are detected simultaneously) mode. In this study, the monocollector mode was used as it obviates the need for complex inter-detector gain calibrations.

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Figure 13. Schematic drawing of the sputtering process. The negative ion beam is causing the emission of secondary ions from the sample surface (after www.cameca.com). Moreover, scanning ion imaging (SII) can be performed with a CAMECA IMS1280. This method allows the determination of element and isotope distribution within a specific area (70 µm x 70 µm) of the sample and is used to study chemical and isotopic variations, such as the distribution of U, Th, and Pb isotopes or Ti (e.g. Harrison and Schmitt 2007, Kusiak et al. 2013, Bellucci et al. 2018). For example, SII can reveal Pb diffusion within single zircon grains in both two and sometimes (Whitehouse et al. 2014) three dimensions. In the case of SII of U-Pb 90 16 238 232 16 238 16 isotopes, the monocollector is used to measure ions of Zr2 O, U, Th O, and U O2, while the multicollector is used to simultaneously measure the distribution of Pb isotopes. SIMS analyses were performed on a CAMECA IMS1280 ion microprobe at the NordSIMS facility, Swedish Museum of Natural History. Data reduction was done using in-house software developed at NordSIMS for the CAMECA IMS1280 analyses and the Excel add-in Isoplot (version 4.15; Ludwig 2012). Ion images were processed using the CAMECA WinImage 2 software.

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6. Summary of manuscripts

6.1 Manuscript I The aim of this study was to investigate whether all of the Apollo 17 breccias formed in one impact event (the Serenitatis impact) or during separate impact events. Previous studies led to the conclusion that all Apollo 17 breccias are a product of the Serenitatis impact event despite the textural differences (Wood 1975, Wolfe et al. 1981). However, additional petrological and chemical analyses showed significant differences between the aphanitic breccias sampled at the South Massif and the micropoikilitic breccias collected at the North Massif. Subsequent studies therefore concluded that these breccias possibly originated from at least two different impact events (Spudis and Ryder 1981).

Four texturally distinct impact breccias were analysed: one subophitic (76015), one aphanitic (72255), and two micropoikilitic breccias (76055, 76215). Rather than just focusing on the geochemical and textural differences, U-Pb ages of Ca-phosphates were obtained in this study to test the previous theories. The samples can be divided into three groups based on a combination of the U-Pb ages and their petrologic relationships: 3920 ± 3 Ma (subophitic), 3922 ± 5 Ma (aphanitic), and 3930 ± 5 Ma (micropoikilitic). In order to get a first order approximation of which impact basins could have formed the breccias, ejecta distribution calculations were used. The results indicate that Nectaris, Crisium, Serenitatis, and Imbrium might be possible candidates.

The age of 3930 ± 5 Ma yielded from the micropoikilitic breccias is interpreted as the age of the Serenitatis impact event, assuming that previous interpretations, which link the micropoikilitic breccias to the Serenitatis basin, are correct. The aphanitic breccia is the only breccia among the four investigated that contains zircon grains, possibly indicating that this breccia originated from within the Procellarum KREEP terrane. Therefore, it might have recorded the Imbrium impact at 3922 ± 5 Ma. If this interpretation is correct, then at least two large impact basins formed within ~8 myr. Additionally, all the other basins interpreted to fall stratigraphically between the Imbrium and Serenitatis impact event must have formed within this narrow time frame of ~8 myr (e.g. Stöffler et al. 2006, Fassett et al. 2012). Another possibility is incomplete resetting of the U- Pb system, which might lead to apparently older ages in the micropoikilitic breccias.

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6.2 Manuscript II Breccia 12013 is composed of two portions: one grey and one dark lithology. The grey portion is mainly composed of felsite and the black portion of lithic fragments of plagioclase and norite. Both parts are thought to have formed in a single impact event. U-Pb SIMS ages of zircon and phosphate grains in breccia 12013 were obtained from both portions of the breccia. The U-Pb ages for the grey portion rage from 4154 ± 7 Ma to 4308 ± 6 Ma (2σ), and for the black portion from 4123 ± 13 Ma to 4328 ± 14 Ma (2σ). Some of the zircon grains located within the grey portion of the breccia show petrological evidence that they grew within the felsite and their ages likely reflect the crystallisation of this lithology. The younger grains exhibit decreasing U and Th concentrations and some grains show recrystallisation features, as well as the formation of neoblasts. The latter usually form at temperatures above 1600-1700 ̊C, which result in decomposition of the primary zircon grain and formation of new zircon as neoblasts. These high temperatures also could have led to open system behaviour of the U-Pb system of the other zircon grains within the felsite, resulting in partial resetting and the observed spread of ages.

The range of ages yielded from zircon in the black portion of the breccia might also be related or incomplete resetting. Nevertheless, these grains occur as fragments within the matrix and have lost the link to their source rocks. The spread of ages could thus also be explained, if the zircon grains kept their original crystallisation ages or if multiple impact events modified the grains. The 207Pb/206Pb age of 3924 ± 3 Ma (95% conf.) obtained from the Ca-phosphate is interpreted as recording the impact event which caused the partial resetting in zircon. The age is also the same, within uncertainties, to the combined average 207Pb/206Pb age of 3926 ± 2 Ma obtained from Ca- phosphate grains within Apollo 14 breccias (Snape et al. 2016), zircon grains in Apollo 12 impact melt breccia (Liu et al 2012), and the SaU 169 (Gnos et al. 2004). This age was interpreted as the Imbrium impact.

The incomplete resetting of the U-Pb system in zircon detected in this study demonstrates the importance to fully understand the processes causing partial Pb loss in order to filter out meaningless and misleading U-Pb ages. While doing so, we will derive a better understanding of the magmatic and impact history of the Moon.

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6.3 Manuscript III U-Pb SIMS ages of four Apollo 14 breccias were obtained, together with a thorough textural and petrological analysis. Several zircon grains grew cogenetically within different lithic clasts. Hence, the U-Pb ages of these grains yielded the primary crystallisation ages of the clasts and magmatic events were identified at ~4286 Ma, ~4200–4220 Ma, and ~4155 Ma. The age distribution of the samples exhibits a minor peak at ~4210 Ma, derived from zircon grains within a noritic clast. This peak is interpreted as true age peak, representing a magmatic event. Another peak is recognized at ~4.33 Ga, which is also present in zircon data from the Apollo 12, 15, and 17 landing sites (e.g. Nemchin et al. 2008, Crow et al. 2017).

An impact event was recorded in a granular zircon grain at 3936 ± 8 Ma. Similarly, a large zircon grain (~430 x 340 µm in size) yielded a weighted average 207Pb/206Pb age of 3941 ± 5 Ma. This grain is interpreted as originating from a KREEP-rich melt, since the magma had to be highly enriched in zirconium to produce this relatively large grain. It either could have formed in an impact melt sheet or it represents late-stage KREEP magmatism. If the latter is the case, the grain provides evidence for a long duration of KREEP magmatism (~400–500 myr; Meyer et al. 1996, Nemchin et al. 2008).

The majority of zircon grains (> 60%) are, however, located within the matrix of the breccias and lack petrological and textural evidence, which could connect them to their original host rocks. The interpretation of their U-Pb ages remains ambiguous and there is a possibility that the U-Pb system of these grains was disturbed and incompletely reset. Therefore, peaks observed in age distribution diagrams might not always represent increased magmatic activity or an impact event. However, the reoccurrence of the same age in distinct samples enhances the confidence that these zircon grains recorded a zircon-forming event, either magmatic or impact related.

7. Conclusions In this thesis, precise U-Pb SIMS ages of Ca-phosphates and zircon from a number of breccia samples collected at the Apollo 12, 14, and 17 landing sites were obtained. Combined with textural and petrological investigation of the samples, discrete impact events and magmatic episodes were identified.

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In particular, the results indicate magmatic events around ~4286 Ma, ~4200–4220 Ma, and ~4155 Ma, as recorded in Apollo 14 zircon grains. These results contradict the simplest interpretation that the majority of zircon grains formed around 4350 Ma and were reset at 3900 Ma. A recurring age in the analysed Ca-phosphates is ~3925 Ma, which is interpreted to be the age of the Imbrium impact (Snape et al. 2016a, Thiessen et al. 2017, 2018). If this interpretation is correct, then Imbrium ejecta indeed covers large regions of the lunar nearside and may have reset the radiometric clocks in many lunar breccias. Nevertheless, Ca-phosphates occurring in three texturally distinctive breccias from the Apollo 17 landing site may have originated from three separate impact events as opposed to a single impact event. Furthermore, slightly older U-Pb ages of ~3940 Ma were recorded in phosphates in Apollo 14 breccias (Snape et al. 2016a) and in two zircon grains analysed in this study. The recurrence of this older age enhances the confidence that these ages date one or several (pre-Imbrium) impact event(s), implying that at least two or even three basin forming impact events occurred within a narrow time interval of ~15 myr. If this scenario is correct, then the Moon experienced a moderate late heavy bombardment, during which several impact basins formed.

This study also demonstrates the importance of conducting thorough textural and petrological investigations of zircon grains, not only to interpret their U-Pb ages as magmatic or impact ages, but also to sort out meaningful U-Pb data from partially reset grains. The observed partial Pb loss in lunar zircon grains is not yet fully understood and requires further investigation. There is a possibility that also some of the Ca-phosphates experienced partial resetting and hence, older phosphates with ages around 3930–3940 Ma (Snape et al. 2016a, Thiessen et al. 2017) might not have recorded a or several (pre-Imbrium) impact event(s).

In conclusion, while this study provides new and precise U-Pb data, which refines the impact and magmatic history of the Moon, this method may have reached its limit to separate impact events within a narrow time interval (< 5 Ma). Taking the results from this study and previously published zircon and Ca-phosphate data (e.g. Nemchin et al. 2009, Merle et al. 2014, Snape et al. 2016a, Crow et al. 2017) into account, the U-Pb data indicate that the Moon experienced a continuous bombardment with several peaks between ~4.5–3.9 Ga, possibly with a moderate spike in impact activity at 3.9 Ga. Future sample return missions will be required to further investigate the early impact flux of the Moon and Solar System.

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8. Acknowledgments I would like to sincerely thank my supervisors Nemchin, Martin Whitehouse, and Victoria Pease for giving me the chance to work on this fascinating and challenging PhD topic. You have been a great help, guidance, and support throughout my PhD and without you I would not be where I am right now.

Many thanks to Josh and Jeremy, who spent a lot of time with me on the SEM and SIMS, corrected manuscripts, and answered questions. I would also like to thank my office and PhD mate Ross for sharing this PhD adventure with me (numerous cakes and ice creams helped us to survive). Furthermore, I would like to thank my colleagues and friends at the NRM for a nice work environment and the cookies and cakes during fika. Leena, you are simply great and helped me from my very first day at NRM. Kerstin, thank you for your help coating the lunar samples. Andreas, thank you for the Swedish translation of the abstract. Melanie (Danke auch für die Korrekturen), Ellen, Lorraine, Cat, Gavin, Renaud: you guys all made a big difference and I will always remember our lunches (outside) or BBQs and swimming at the lake.

I would like to say thank you to my PhD colleagues at SU, who made this PhD journey a lot of fun. Many thanks to my Stockholm friends outside the SU for fun nights, dinners, watching German football matches, many fika: Alena, Isabel, Franz, Duarte, Cécile, Fiona, Gaby. Alex, thank you for being a lovely confused Frenchman. You are a good friend! Pedro, you are just awesome and a great friend.

I also would like to thank the Apollo astronauts, who risked their lives collecting the samples that I have investigated. It is still one of my dreams to travel into space myself and you guys are a great inspiration and role model for future generations!

To all my friends around the world: I miss you guys and your friendships are very precious to me. Sanda, Markus, Steffi, David (reinhaun!), Marie, Evi and Daniel, Sami, Ines, Dominik, Lars, my LEAPS friends, my Couchsurfing acquaintances: you are all wonderful! Birte, unsere jahrelange Freundschaft ist etwas ganz Besonderes und ich werde niemals unsere Abenteuer in Australien vergessen. Vicky & Thilo, unsere Trips nach Gotland, Barcelona oder einfach nur Partyabend in MS sind immer großartig! Julia, ich bewundere deine Stärke und es ist toll, dass wir unsere Leidenschaft für die Planetologie teilen. Anne, ich bin sehr froh, dass sich unsere Wege durchs

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Studium gekreuzt haben und wir so gute Freunde geworden sind. Ich kann immer auf dich zählen. Pia, thanks for your craziness and laughter.

I also would like to thank my Dutch family and friends for great times in Hoogeveen. I felt at home from the very first day and the Netherlands have become a big part of me.

My deep gratitude goes to my wonderful parents, who supported me throughout my entire life – even when I decided that I wanted to try to become an astronaut. Danke für eure Liebe und dass ihr immer für mich da seid. Ohne euch wäre ich nie so weit gekommen. Weiterhin geht ein großes Dankeschön an meine Schwester Laura. Ich denke immer gerne zurück an unser unsere gemeinsamen Reisen nach Australien, Kuba, Jamaika, Fidschi etc. und an wilde (Dorf-)Partys. Ein riesiges Dankeschön auch an meine Großeltern für ihre Liebe und Unterstützung. Ich hätte mir keine besseren Großeltern wünschen können und ihr seid für immer in meinem Herzen.

Last but not least, I would like to express my deep gratitude to my partner Jan Jaap. Thank you for all the sacrifices you have made for me. You are a great support, my best travel mate, and a wonderful father. The world would not be the same anymore without our beautiful son Finn, who teaches me a lot every day and is making me extremely happy. Du bist das Beste, was mir je passiert ist. Ik hou heel veel van jullie.

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