Encyclopedia of Lunar Science 3319145401, 9783319145402

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Brian Cudnik Editor

Encyclopedia of Lunar Science

Encyclopedia of Lunar Science

Brian Cudnik Editor

Encyclopedia of Lunar Science With 641 Figures and 79 Tables

Editor Brian Cudnik Prairie View A&M University Houston, TX, USA

ISBN 978-3-319-14540-2 ISBN 978-3-319-14541-9 (eBook) https://doi.org/10.1007/978-3-319-14541-9 © Springer Nature Switzerland AG 2023 This work is subject to copyright. All rights are reserved by the Publisher, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilms or in any other physical way, and transmission or information storage and retrieval, electronic adaptation, computer software, or by similar or dissimilar methodology now known or hereafter developed. The use of general descriptive names, registered names, trademarks, service marks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. The publisher, the authors, and the editors are safe to assume that the advice and information in this book are believed to be true and accurate at the date of publication. Neither the publisher nor the authors or the editors give a warranty, expressed or implied, with respect to the material contained herein or for any errors or omissions that may have been made. The publisher remains neutral with regard to jurisdictional claims in published maps and institutional affiliations. This Springer imprint is published by the registered company Springer Nature Switzerland AG. The registered company address is: Gewerbestrasse 11, 6330 Cham, Switzerland

List of Topics

Character and Spatial Distribution of Materials and Structures on the Surface of the Moon Section Editors: Yuanzhi Zhang and Shengbo Chen Chronology, Lunar Samples Early Lunar Structures, Engineering Iron Concentration, Lunar Surface Lunar Elemental Distribution Lunar Geological Timescale Lunar Landscape, Highlands Lunar Landscape, Impact Craters Lunar Landscape, Maria Lunar Landscape: Domes Lunar Landscape: Grabens Lunar Landscape: Rilles Lunar Landscape: Wrinkle Ridges Lunar Lava Tubes Lunar Magma Ocean Lunar Meteorites Lunar Mineral Distribution Lunar Permanently Shaded Areas Lunar Regolith: Materials Lunar Rocks Lunar Surface, Bulk Density and Porosity Lunar Surface, Dielectric Permittivity Lunar Surface, Electrical Conductivity Lunar Surface, Gravity Field Lunar Surface, Magnetic Field Lunar Surface, Seismic Properties Lunar Terrane Tectonics Reflection and Emission of Radiation from the Moon

Reflectivity Regolith Physical Properties Regolith Structure Surface of the Moon, Distribution of Materials and Structures

Formation and Crystallization of the LMO Section Editor: Jesse Davenport Evolution, Lunar: From Magma Ocean to Crust Formation Lunar Magma Ocean Theory, Origins, and Rationale Lunar Magma Ocean, A Benchmark of Lunar and Planetary Geologic Theory Lunar Magma Ocean, Comparison to Other Planetary Magma Oceans Lunar Magma Ocean, Composition of the Bulk LMO Lunar Magma Ocean, Pre-Apollo, Apollo, and Post-Apollo Views Lunar Magma Ocean, Size Water in the LMO

Geologic and Impact Processes on the Moon’s Surface Section Editors: Yuanzhi Zhang and Dr. Jianzhong Liu Apollo 17: The Last Time Humans Walked on the Moon Basalt v

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Breccia Color of the Moon Early Bombardment Impact Ejecta History of Lunar Science I History of Lunar Science II Moon Optical Maturity (OMAT) Weathering

List of Topics

Lunar Interior, Geophysical Models Lunar Interior, Halogens Lunar Melts, Densities of Lunar Melts, Viscosity of Lunar Primitive Crust, Evolution of Mantle Convection Origin and Evolution of the Moon: Tungsten Isotopic Constraints Siderophile Elements in the Lunar Mantle Sulfides in the Moon

Lunar Atmosphere Section Editor: Cesare Grava Contribution of Surface Processes to the Lunar Exosphere Lunar Atmosphere Lunar Atmosphere, Alkali Lunar Exosphere Lunar Atmosphere, Composition Lunar Atmosphere, Effects of Cometary Impacts Lunar Atmosphere, Energetic Neutral Atoms Lunar Atmosphere, Source and Loss Processes Lunar Atmosphere, Transport and Storage of Volatiles Lunar Dust Exosphere Lunar Ionosphere Noble Gases

Lunar Impact Flashes Section Editor: José M. Madiedo Impact Event, Total Lunar Eclipse Lunar Impact Event: The 11 September 2013 Lunar Impact Events, Luminous Efficiency, and Energy of Lunar Impact Flashes, Causes and Detection Lunar Impact Flashes, Temperature MIDAS System

Lunar Missions Section Editor: Dr. Caitlin Ahrens Apollo Lunar Roving Vehicles Apollo Program ARTEMIS Mission Chandrayaan-1 Mission Chandrayaan-2 Mission Chang’e Missions Cis-lunar Trajectory Clementine Mission GRAIL Mission Hiten Mission Kaguya (SELENE) Mission LRO Mini-RF Luna Missions Lunar Prospector Mission Lunar Reconnaissance Orbiter (LRO) Mission Ranger Missions SMART-1 Mission Surveyor: The Science Surveyor: The Spacecraft The Lunar Atmosphere and Dust Environment Experiment (LADEE) Mission Zond Missions

Mineralogy, Petrology and Geochemistry of the surface of the Moon Lunar Interior Section Editor: Edgar Sikko Steenstra Lunar Core Composition Lunar Core Dynamo Lunar Core Formation

Section Editor: Dr. Amit Basu Sarbadhikari Cryptomare and Its Mineralogy Regolith Thickness Silica, Silicate Young Volcanism on the Moon

List of Topics

Observation and Data Collection Techniques Section Editors: Dr. Prakash Chauhan and Dr. Mamta Chauhan Infrared Spectroscopy Legacy of the Apollo Seismic Experiments Lunar Ultraviolet Spectroscopy Magnetometry Polar Ice on the Moon Polarimetry Measurements of the Moon Pyroclastic Deposits, Remote Sensing of Radio Astronomy with the Moon Reflectance Spectroscopy Remote Sensing in the Thermal Infrared Selenodesy Visible and Near-Infrared Spectroscopy

Orbital Dynamics and Tidal Interactions Section Editor: Dr. Paul M. Bremner Accretion of the Moon Barycenter of the Earth-Moon System Ellipsoid Gravitational Field Measurements to Deduce the Lunar Interior Structure Laser Ranging Lunar Libration Moon: Seismicity Saros Cycle Satellite, Natural Sidereal Period

Other Topics in Lunar Science Section Editors: Nicolle E. B. Zellner and Dr. Caitlin Ahrens Crystallization of the Lunar Magma Ocean Detection of Water Early Geologic History of the Moon Estimate of Lunar TiO2 and FeO with M3 Data

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Impact Processes on the Moon Internal Structure/Mantle Motions of the Moon Isotopic Composition of the Moon and the Lunar Isotopic Crisis Laboratory Analysis (Reflectance Spectroscopy) of Terrestrial Analogues Large-Scale Faulting on the Moon LCROSS, Lunar Diviner Instrument Lunar Cartography Lunar Crater Ejecta Lunar Crater Observation and Sensing Satellite (LCROSS) Lunar Domes Lunar Dust Toxicity Lunar Impact Melt Deposits Lunar Length of Day Lunar Magnetic Anomalies Lunar Mare Basalts, Stratigraphy of Lunar Photometry Lunar Regolith Simulants Lunar Surface Composition from X-ray Fluorescence Measurements Lunar Surface, Interaction of the Solar Wind with Upper Regolith Lunar Tectonics Lunar Tectonism, History of Lunar Transient Phenomena Mantle Mascons Modeling of the Lunar Magma Ocean Moon: Origin, Alternative Theories Origin of the Moon, Impactor Theory Plasma Environment of the Moon Regolith Production by Thermal Stress Weathering Silicic Volcanism on the Moon Subsurface Geology from Remote Sensing Surface and Near-Surface Thermal Environment of the Moon Surveyor Imagery Thermal Stress Weathering of Boulders on Airless Bodies Topographic Studies of the Moon Volatiles on the Lunar Surface and Subsurface Volcanic Processes on the Moon

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Physics and Physical Processes of the Moon Section Editor: Dr. Caitlin Ahrens Absorption, Absorption Spectrum in Lunar Studies Albedo of Lunar Features Center of Mass and Center of Figure of Moon, Gravity, and Inertia Electromagnetic Radiation Inertia, Inertial Frame Magnetic Properties at the Lunar Surface Opposition Effect Shock Metamorphism

Radiation Environment and Interactions Section Editor: Dr. Paul M. Bremner Cosmic Rays in the Lunar Environment Lunar Occultations, Grazing

List of Topics

Radiation Environment of the Moon Total Lunar Occultations Ultraviolet Radiation

Thermal Properties of the Surface and Interior Section Editor: Dr. K. Durga Prasad Differentiation of the Lunar Interior Early Thermal Evolution of the Lunar Interiors Eclipse, Lunar Eclipse, Solar Emissivity Infrared Radiation Insolation and Temperature on the Moon Lunar Crust, Chemical Composition Lunar Crust, Morphology Thermophysical Behavior of the Lunar Surface

About the Editor

Professor Brian Cudnik serves as Laboratory Coordinator II and Adjunct Instructor for the Physics Program at Prairie View A&M University (PVAMU, a part of the Texas A&M University system) in Texas. He has been in this position for over 21 years and has been at Prairie View A&M for a total of 24 years. His prior position at PVAMU was Research Assistant at the Solar Observatory. He has served as coordinator of the Lunar Meteoritic Impact Search section of the Association of Lunar and Planetary Observers (ALPO) since January 2000, 2 months after making the first scientifically confirmed visual observation of a meteoroid impact on the Moon during the Leonid storm of November 1999.He has published papers and posters on various astronomical subjects including peer-reviewed papers, posters at professional conferences, and amateur astronomy publications. Professor Cudnik has also served as Board Member of the Houston Astronomical Society, is presently a Full member of the American Astronomical Society, and is a member of the American Association of Variable Star Observers, to which he regularly contributes observations. Additionally, he has taught astronomy at the University of St. Thomas every semester; including summer sessions (one summer session per summer) from 2005 to 2015.

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Associate Editor

Caitlin Ahrens NASA Goddard Space Flight Center Greenbelt, MD, USA

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Section Editors

Caitlin Ahrens NASA Goddard Space Flight Center, Greenbelt, MD, USA Amit Basu Sarbadhikari Physical Research Laboratory, Ahmedabad, India Paul M. Bremner Heliophysics and Planetary Science Branch, NASA, Huntsville, AL, USA Mamta Chauhan Geosciences Division, Indian Institute of Remote Sensing (IIRS), Indian Space Research Organization (ISRO), Dehradun, Uttarakhand, India Prakash Chauhan National Remote Sensing Centre (NRSC), Indian Space Research Organization (ISRO), Hyderabad, Telangana, India Shengbo Chen College of Geo-exploration Science and Technology, Jilin University, Changchun, China Jesse Davenport DeLand, FL, USA K. Durga Prasad Planetary Sciences Division, Physical Research Laboratory, Ahmedabad, India Cesare Grava Southwest Research Institute, San Antonio, TX, USA Jianzhong Liu Center for Lunar and Planetary Sciences, Institute of Geochemistry, Chinese Academy of Sciences, Guiyang, China José M. Madiedo Solar System Department, Instituto de Astrofísica de Andalucía, CSIC, Granada, Spain Edgar Sikko Steenstra Faculty of Earth and Life Sciences, Vrije Universiteit, Amsterdam, The Netherland Nicolle E. B. Zellner Department of Physics, Albion College, Albion, MI, USA Yuanzhi Zhang Key Laboratory of Lunar and Deep-space Exploration, Chinese Academy of Science, Beijing, PR, China

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Contributors

Caitlin Ahrens NASA Goddard Space Flight Center, Greenbelt, MD, USA Sami W. Asmar NASA’s Jet Propulsion Laboratory, California Institute of Technology, Pasadena, CA, USA Wayne Bailey Association of Lunar and Planetary Observers, Sewell, NJ, USA Edward Baker Earth Science, University of Oxford, Oxford, UK Donald Barker MAXD, Inc., Houston, TX, USA Thomas J. Barrett School of Physical Sciences, The Open University, Milton Keynes, UK Amit Basu Sarbadhikari Physical Research Laboratory, Ahmedabad, India Jeffrey Baumgardner Center for Space Physics, Boston University, Boston, MA, USA James F. Bell III Arizona State University, Tempe, AZ, USA Haym Benaroya Rutgers University, Piscataway, NJ, USA Kristen Bennett Northern Arizona University, Flagstaff, AZ, USA Leonhard E. Bernold Universidad Tecnica Federico Santa Maria, Valparaiso, Chile Rajneesh Bhutani Department of Earth Sciences, Pondicherry University, Puducherry, India Doris Breuer German Aerospace Center (DLR), Institute of Planetary Research, Berlin, Germany Christoph Burkhardt Institut für Planetologie, Westfälische WilhelmsUniversität Münster, Münster, Germany Sönke Burmeister Department of Physics, Kiel University, Kiel, Germany Paul K. Byrne Planetary Research Group, Department of Marine, Earth, and Atmospheric Sciences, North Carolina State University, Raleigh, NC, USA Om Prakash Narayan Calla International Centre for Radio Science (ICRS), Jodhpur, India xv

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C. Carli Inaf-IAPS, Tor Vergata, Rome, Italy Julien Chabé Université Côte d’Azur, Observatoire de la Côte d’Azur, CNRS, IRD, Géoazur, Caussols, France Mamta Chauhan Department of Geology, School of Earth Sciences, Banasthali Vidyapith, Niwai, Rajasthan, India Geosciences Division, Indian Institute of Remote Sensing (IIRS), Indian Space Research Organization (ISRO), Dehradun, Uttarakhand, India Prakash Chauhan Indian Institute of Remote Sensing (IIRS), Indian Space Research Organization (ISRO), Dehradun, Uttarakhand, India National Remote Sensing Centre (NRSC), Indian Space Research Organization (ISRO), Hyderabad, Telangana, India Bradley Cheetham Advanced Space, LLC, Boulder, CO, USA Shengbo Chen College of Geo-exploration Science and Technology, Jilin University, Changchun, Jilin, China Nicholas Connors Department of Physics, McGill University, Montréal, QC, Canada Anthony Cook Department of Physics, University of Aberystwyth, Aberystwyth, Ceredigion, UK Ron Creel Huntsville, AL, USA Francis A. Cucinotta Univerity of Nevada Las Vegas, Las Vegas, NV, USA Brian Cudnik Department of Chemistry and Physics, Prairie View A&M University, Prairie View, TX, USA Jesse Davenport Cactus Communications, Deland, Florida, USA Liu Dawei National Astronomical Observatories, Chinese Academy of Sciences, Beijing, China Jan Deca Laboratory for Atmospheric and Space Physics, University of Colorado Boulder, Boulder, CO, USA Deepak Dhingra Department of Earth Sciences, Indian Institute of Technology Kanpur, Uttar Pradesh, India Min Ding School of Earth and Space Sciences, Peking University, Beijing, China Patrick Donohue Hawai’i Institute of Geophysics and Planetology, University of Hawai’i at Manoa, Honolulu, HI, USA Catherine A. Dukes Laboratory for Astrophysics and Surface Physics, University of Virginia, Charlottesville, VA, USA K. Durga Prasad Planetary Sciences Division, Physical Research Laboratory, Ahmedabad, India

Contributors

Contributors

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Kenneth L. Edmundson Edmundson Photogrammetry Consulting LLC, Flagstaff, AZ, USA Jennifer E. Edmunson Jacobs Technologies, Inc., Huntsville, AL, USA Vincent Eke Institute for Computational Cosmology, Durham University, Durham, UK Stephen Elardo Geophysical Laboratory, Carnegie Institution of Washington, Washington, DC, USA R. C. Elphic Planetary Systems Branch, MS 245-3, NASA Ames Research Center, Moffett Field, CA, USA Kimberly Ennico-Smith NASA Ames Research Center, Moffett Field, CA, USA Alexander J. Evans Lunar and Planetary Laboratory, University of Arizona, Tucson, AZ, USA Wenzhe Fa Institute of Remote Sensing and Geographical Information System, Peking University, Beijing, China Amy L. Fagan Geosciences and Natural Resources Deptartment, Western Carolina University, Cullowhee, NC, USA William H. Farrand Space Science Institute, Boulder, CO, USA Bernard H. Foing ESA – European Space Research and Technology Centre, Noordwijk, The Netherlands ILEWG EuroMoonMars, VU Amsterdam & Leiden Observatory, Amsterdam, The Netherlands Rahul Dev Garg Geomatics Engineering group, Department of Civil Engineering, Indian Institute of Technology, Roorkee, Roorkee, Uttarakhand, India Aleksandra J. Gawronska Department of Geology and Environmental Earth Science, Miami University, Oxford, OH, USA Robin H. Glefke UC San Diego, San Diego, CA, USA David B. Goldstein The University of Texas at Austin, Austin, TX, USA Vishal Goyal Department of Physics, Panjab University, Chandigarh, India Cesare Grava Southwest Research Institute, San Antonio, TX, USA Juliane Gross Department of Earth and Planetary Sciences, Rutgers University, Piscataway, NJ, USA American Museum of Natural History, New York, NY, USA Lunar and Planetary Institute, Houston, TX, USA Dijun Guo Center for Lunar and Planetary Sciences, Institute of Geochemistry, Chinese Academy of Sciences, Guiyang, Guizhou, China Saira Hamid School of Earth and Space Exploration, Arizona State University, Tempe, AZ, USA

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Bruce Hapke Department of Geology and Environmental Science, University of Pittsburgh, Pittsburgh, PA, USA Tara S. Hayden School of Physical Sciences, The Open University, Milton Keynes, UK Harald Hiesinger Institut für Planetologie, Westfälische WilhelmsUniversität, Münster, Germany Lon L. Hood Lunar and Planetary Laboratory University of Arizona, Tucson, AZ, USA Mihaly Horanyi Laboratory for Atmospheric and Space Physics, and Department of Physics, University of Colorado, Boulder, CO, USA Gordon L. Houston American Astronomical Society, Houston, TX, USA Shuang Huang College of Geo-exploration Science and Technology, Jilin University, Changchun, China Zhaojun Huang The Sino-Forest Applied Research Centre for Pearl River Delta Environment, Hong Kong Baptist University, Kowloon Tong, Hong Kong, China Jinzhu Ji Center for Lunar and Planetary Sciences, Institute of Geochemistry, Chinese Academy of Sciences, Guiyang, China Ping Jinsong National Astronomical Observatories of CAS, Key Laboratory of Lunar and Deep-space Exploration, Chinese Academy of Sciences, Beijing, China Robert E. Johnson University of Virginia, Charlottesville, VA, USA Katherine H. Joy Department of Earth and Environmental Sciences, The University of Manchester, Manchester, UK P. Kalyana Srinivasa Reddy Planetary Sciences Division, Physical Research Laboratory, Ahmedabad, India Jacob Kegerreis Institute for Computational Cosmology, Durham University, Durham, UK Akos Kereszturi Research Centre for Astronomy and Earth Sciences, Konkoly Thege Miklos Astronomical Institute, Budapest, Hungary Walter S. Kiefer Lunar and Planetary Institute/USRA, Houston, TX, USA Rosemary M. Killen NASA Goddard Space Flight Center, Greenbelt, MD, USA Thorsten Kleine Institut für Planetologie, University of Münster, Münster, Germany G. G. Kochemasov IGEM of the Russian Academy of Sciences, Moscow, Russian Federation

Contributors

Contributors

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Thomas S. Kruijer Institut für Planetologie, University of Münster, Münster, Germany Nuclear and Chemical Sciences Division, Lawrence Livermore National Laboratory, Livermore, CA, USA K. N. Kusuma Department of Earth Sciences, Pondicherry University, Pondicherry, India David J. Lawrence Johns Hopkins University Applied Physics Laboratory, Laurel, MD, USA Raffaello Lena Geologic Lunar Research (GLR) Group, Rome, Italy Chunlai Li National Astronomical Observatory, Chinese Academy of Sciences, Beijing, China Key Laboratory of Lunar and Deep Space Exploration, National Astronomical Observatories, Chinese Academy of Sciences, Beijing, China Yi-Li Lin McGill University, Montréal, QC, Canada Bin Liu National Astronomical Observatories, Chinese Academy of Sciences, Beijing, China Key Laboratory of Lunar and Deep-Space Exploration, Chinese Academy of Science, Beijing, China Jianzhong Liu Center for Lunar and Planetary Sciences, Institute of Geochemistry, Chinese Academy of Sciences, Guiyang, Guizhou, China Martin J. Losekamm Department of Physics, Technical University of Munich, Garching, Germany Excellence Cluster ORIGINS, Garching, Germany Ming Ma College of Geo-exploration Science and Technology, Jilin University, Changchun, Jilin, China José M. Madiedo Solar System Department, Instituto de Astrofísica de Andalucía, CSIC, Granada, Spain Facultad de Ciencias Experimentales, Universidad de Huelva, Huelva, Spain Samuel W. (Walt) McCandless Hughes Aircraft, Surveryor Project (19601969), Tucson, AZ, USA Claire McLeod Department of Geology and Environmental Earth Science, Miami University, Oxford, OH, USA Zhiguo Meng College of Geo-exploration Science and Technology, Jilin University, Changchun, China Bernard P. Miller RCA Ranger Project, Princeton, NJ, USA Kaushik Mitra Department of Geosciences, Stony Brook University, Stony Brook, NY, USA Donald R. Montgomery Jet Propulsion Laboratory, CalTech, Surveyor Project (1962-1966), Ashland, OR, USA

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Shyama Narendranath Indian Space Research Organization, Bangalore, India Richard Nugent International Occultation Timing Association, Dripping Springs, TX, USA Ceri Nunn Jet Propulsion Laboratory, California Institute of Technology, Pasadena, CA, USA Deniz Ölçek University of Oslo, Center for Space Sensors and Systems, Oslo, Norway Joseph G. O’Rourke School of Earth and Space Exploration, Arizona State University, Tempe, AZ, USA José L. Ortiz Instituto de Astrofísica de Andalucía, CSIC, Granada, Spain Gordon R. Osinski Department of Earth Sciences, Institute for Earth and Space Exploration, University of Western Ontario, London, ON, Canada Ziyuan Ouyang National Astronomical Observatory, Chinese Academy of Sciences, Beijing, China Key Laboratory of Lunar and Deep Space Exploration, National Astronomical Observatories, Chinese Academy of Sciences, Beijing, China Jinsong Ping National Astronomical Observatory of CAS, Key Laboratory of Lunar and Deep-space Exploration, Chinese Academy of Sciences, Beijing, China Nicola J. Potts School of GeoSciences, University of Edinburgh, Edinburgh, UK K. Durga Prasad Planetary Sciences Division, Physical Research Laboratory, Ahmedabad, India Parvathy Prem The University of Texas at Austin, Austin, TX, USA Giuseppe D. Racca ESA – European Space Research and Technology Centre, Noordwijk, The Netherlands Nachiketa Rai Department of Earth Sciences, Indian Institute of Technology, Roorkee, Roorkee, India V. J. Rajesh Department of Earth and Space Sciences, Indian Institute of Space Science and Technology (IIST), Thiruvananthapuram, Kerala, India Jon Rask Space Biosciences Research Branch, NASA Ames Research Center, KBRWyle, Moffett Field, CA, USA Justin J. Rennilson Jet Propulsion Laboratory, CalTech, Surveyor Project (1961-1969), La Mesa, CA, USA Kurt D. Retherford Southwest Research Institute, San Antonio, TX, USA Aaron J. Rosengren University of California San Diego, Ja Jolla, CA, USA Vijayan S. PLANEX, Physical Research Laboratory, Ahmedabad, India

Contributors

Contributors

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Premkumar B. Saganti Prairie View A&M, Prairie View, TX, USA Sandeep Sahijpal Department of Physics, Panjab University, Chandigarh, India Daniel J. Scheeres University of Colorado Boulder, Boulder, CO, USA Carl Schmidt Center for Space Physics, Boston University, Boston, MA, USA Richard W. Schmude Jr. Gordon State College, Barnesville, GA, USA Norbert Schörghofer Institute for Astronomy, University of Hawaii at Manoa, Honolulu, HI, USA Elliot Sefton-Nash Department of Earth and Planetary Sciences, Birkbeck, University of London, London, UK Jayanth Koundinya Serla School of Earth and Space Exploration, Arizona State University, Tempe, AZ, USA G. Serventi Department of Physics and Earth Sciences “Macedonio Melloni”, University of Parma, Parma, Italy M. Sgavetti Department of Physics and Earth Sciences “Macedonio Melloni”, University of Parma, Parma, Italy N. Srivastava Planetary Sciences & Exploration Programme, Physical Research Laboratory, Ahmedabad, India Yash Srivastava Physical Research Laboratory, Ahmedabad, India Indian Institute of Technology Gandhinagar, Gandhinagar, India Edgar Sikko Steenstra Faculty of Earth and Life Sciences, Vrije Universiteit, Amsterdam, The Netherlands Philip J. Stooke Department of Geography, University of Western Ontario, London, ON, Canada B. E. Strauss National Institute of Standards and Technology, Gaithersburg, MD, USA Timothy J. Stubbs Solar System Exploration Division, NASA Goddard Space Flight Center, Greenbelt, MD, USA Jianing Su McGill University, Montreal, QC, Canada Ying Sun College of Geoexploration Science and Technology, Jilin University, Changchun, China Jamey R. Szalay Department of Astrophysical Sciences, Princeton University, Princeton, NJ, USA Romain Tartèse Muséum National d’Histoire Naturelle, Institut de Minéralogie, de Physique des Matériaux et de Cosmochimie (IMPMC), Paris, France

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Vivek Thaker Department of Physics and Computer Science, McGill University, Montréal, QC, Canada P. M. Thesniya Department of Earth and Space Sciences, Indian Institute of Space Science and Technology (IIST), Thiruvananthapuram, Kerala, India Jean-Marie Torre Université Côte d’Azur, Observatoire de la Côte d’Azur, CNRS, IRD, Géoazur, Caussols, France Prateek Tripathi Geomatics Engineering group, Department of Civil Engineering, Indian Institute of Technology, Roorkee, Roorkee, Uttarakhand, India Indhu Varatharajan Institute of Planetary Research, German Aerospace Center DLR, Berlin, Germany Department of Geosciences, Stony Brook University, New York, NY, USA Jan Viatteau Faculty of Physics, McGill University, Montreal, Canada A. Vorburger University of Bern, Bern, Switzerland Haley Wahl West Virginia University, Morgantown, WV, USA D. Waller Johns Hopkins Applied Physics Laboratory, Laurel, MD, USA Xianmin Wang Hubei Subsurface Multi-scale Imaging Key Laboratory, Institute of Geophysics and Geomatics, China University of Geosciences, Wuhan, China Tristram Warren Atmospheric, Oceanic and Planetary Physics, University of Oxford, Oxford, UK Wim van Westrenen Faculty of Earth and Life Sciences, Vrije Universiteit, Amsterdam, The Netherlands Richard P. Wilds Division for Planetary Sciences/Historical Astronomy Division, American Astronomical Society, Williamsville, NY, USA Ke Wu Hubei Subsurface Multi-scale Imaging Key Laboratory, Institute of Geophysics and Geomatics, China University of Geosciences, Wuhan, China P. Wurz University of Bern, Bern, Switzerland Tianqi Xie Department of Earth Sciences/Institute for Earth and Space Exploration, University of Western Ontario, London, ON, Canada Hongwei Yang Chinese Academy of Geological Sciences, Beijing, China Michael Zanetti University of Western Ontario, London, ON, Canada Nicolle E. B. Zellner Department of Physics, Albion College, Albion, MI, USA National Astronomical Observatories, Chinese Academy of Sciences, Beijing, China Guangliang Zhang National Astronomical Observatory, Chinese Academy of Sciences, Beijing, China

Contributors

Contributors

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Key Laboratory of Lunar and Deep Space Exploration, National Astronomical Observatories, Chinese Academy of Sciences, Beijing, China Nan Zhang Department of Geological Science, Lunar and Planetary Institute, Brown University, Providence, RI, USA Weijia Zhang Department of Physics, University of Oxford, Oxford, UK Liang Zhao College of Geo-exploration Science and Technology, Jilin University, Changchun, China Wenjin Zhao Chinese Academy of Geological Sciences, Beijing, China Yongchun Zheng National Astronomical Observatory, Chinese Academy of Sciences, Beijing, China Key Laboratory of Lunar and Deep Space Exploration, National Astronomical Observatories, Chinese Academy of Sciences, Beijing, China Meng Zhiguo College of Geo-exploration Science and Technology, Jilin University, Changchun, China Chao Zhou College of Geo-exploration Science and Technology, Jilin University, Changchun, Jilin, China National Marine Environmental Monitoring Center, Dalian, Liaoning, China Qin Zhou Key Laboratory of Lunar and Deep Space Exploration, Chinese Academy of Sciences, Beijing, China National Astronomical Observatories, Chinese Academy of Sciences, Beijing, China Yongliao Zou National Astronomical Observatory, Chinese Academy of Sciences, Beijing, China Key Laboratory of Lunar and Deep Space Exploration, National Astronomical Observatories, Chinese Academy of Sciences, Beijing, China

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Absorption, Absorption Spectrum in Lunar Studies Prateek Tripathi and Rahul Dev Garg Geomatics Engineering Group, Department of Civil Engineering, Indian Institute of Technology, Roorkee, Roorkee, India

Definition Absorption of solar radiation by the rocks on the lunar surface plays a crucial role in identifying a particular mineral’s spectral absorption feature in the electromagnetic spectrum (EMS). Spectroscopy, which is majorly based on the “absorption” caused by the interaction of solar electromagnetic wave radiation with the rocks and soils on the lunar surface, pertains to two primary goals, first is to study the matter structures and then internal interactions (e.g., atomic and molecular structures along with various orbital, spin, and nuclear interactions) followed by the identification of chemical composition and quantity measurement. However, the “interaction” and “absorption” significantly modify depending on the particular range of wavelength in EMS (Clark 1999). Optical spectroscopy utilizes the visible to near-infrared and shortwave infrared (VNIR-SWIR) region (400–3000 nm) (metal-to-metal charge transfer transitions) and in the far-infrared range (3000–10,000 nm) where excited state has the option of being in the present state and emitting © Springer Nature Switzerland AG 2023 B. Cudnik (ed.), Encyclopedia of Lunar Science, https://doi.org/10.1007/978-3-319-14541-9

a photon between filled and unfilled 3d orbitals, later resulting in motions (Fig. 1). Also, transfer of charge between oxygen and metal results in absorption of energy in the near-ultraviolet (oxygen-to-metal charge transfer transitions) (Adams and Jones 1970; Adams 1974). In absorption concept for minerals, many effects are possible, including the charge transfer effect, the conduction-band absorption effect, electronic transition in transition metals, and crystal field effect. The absorbed energy has a different value for different mineral species. Major molecules/ ions responsible for electronic motions are Fe, P-O, Al-O (Clark 1999; Chukanov and Chervonnyi 2016). The atomic-molecular vibrational motions in the SWIR and TIR (thermal infrared) regions of the EMS are the major reason for identifying rockforming minerals, including carbonates, silicates, hydroxyls, oxides, etc. (Burke 1986; Schrader 1996; Griffiths and Chalmers 2002; MARTENS 2004). These processes arise due to stretching and bending motions, and can be identified as fundamental tones mainly in the thermal infrared region (> 3500 nm) and their combinations and overtones in the SWIR region (1000–3000 nm) (Gupta 2003).

Lunar Spectroscopy and Sensors The solar radiation reflected from the lunar surface helps identify the four most common lunar

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Absorption, Absorption Spectrum in Lunar Studies

Absorption, Absorption Spectrum in Lunar Studies, Fig. 1 The physical phenomenon of the interaction of radiation on the lunar surface. The reflected and emitted light gets captured by the sensor on the spacecraft. The phenomenon of photon transfers and electronic transition

takes place. In the last, the reflectance spectra in the VNIRSWIR range are obtained with two major absorption features near 900 nm and 1500 nm. The spectra shown are from the Shoemaker impact crater (88
60 S, 44
540 E)

minerals, i.e., olivine, plagioclase, pyroxene, and ilmenite, each of which has a unique footprint (Dhingra 2018). These minerals and their absorption features were recorded with imaging and nonimaging spectrometers on an orbiting spacecraft launched by different space agencies. Following the results from the spectral analysis of lunar rocks and soils returned from Apollo missions, several countries had launched the orbital and robotic missions with imaging and non-imaging spectrometers as payload, to study the lunar surface properties in different wavelength ranges of EMS. These spectrometers are either multispectral or hyperspectral and work in ultraviolet, visible to infrared, short wave infrared, and thermal infrared range. The primary imaging and non-imaging spectrometers are: Ultraviolet/Visible camera, Near-Infrared CCD Camera, and Longwave Infrared (LWIR) Camera onboard Clementine, a joint mission of DOD/NASA (USA) launched in 1994 (Nozette et al. 1994), Infrared Spectrometer (SIR) onboard Small Missions for Advanced Research in Technology-1 (SMART-1) of European Space Agency launched in 2003 (Foing et al. 2003), the Interference Imaging spectrometer, Infrared

spectrometer, VNIS (Visible and Near-Infrared Imaging Spectrometer) and Lunar Mineral Spectrometer (LMS) on aboard Chang’e series (2004-present) by China National Space Administration (CNSA)(He et al. 2014; Ling et al. 2019), Multiband Imager (MI) and Spectral profiler (SP) onboard SELenological and ENgineering Explorer “KAGUYA” (SELENE) of Japan Aerospace Exploration Agency (JAXA) launched in 2007 (Ohtake et al. 2008), HySI (Hyper Spectral Imager), Moon Mineralogy Mapper (M3), SIR-2 (Near-infrared spectrometer) on Chandrayaan-1 (2008) of Indian Space Research Organisation (ISRO) (Anbazhagan and Arivazhagan 2009; Pieters et al. 2009), Diviner Lunar Radiometer Experiment (DLRE) on board Lunar Reconnaissance Orbiter (LRO) and Lunar Crater Observation and Sensing Satellite (LCROSS) by National Aeronautics and Space Administration (NASA) in 2009 (Robinson et al. 2010), UV-Vis Spectrometer (UVS) onboard Lunar Atmosphere and Dust Environment Explorer (LADEE) launched in 2013 (Horanyi et al. 2014), and Imaging Infrared Spectrometer (IIRS) onboard Chandrayaan-2 (2019) (Chauhan et al. 2021).

Absorption, Absorption Spectrum in Lunar Studies

Lunar Geology and Its Impact on Remote Absorption Profiles However, the well-known technique used to identify the absorption on the lunar surface is reflectance spectroscopy because (1) the intensities of the incoming solar radiation in the wavelength range of 300–1000 nm are highly variable; (2) the lunar surface mineralogy has much variety in composition. One of the two major regions on the moon, Highlands, the bright areas, are entirely composed of anorthosites, which are rich in aluminum, calcium, and extremely widespread KREEP rocks enriched in K (potassium), P (phosphorous), and REE (rare earth element). Anorthosites have calcium-rich plagioclase and low calcium pyroxene (in significantly less amount), olivine, clinopyroxene, and Mg-rich rocks, which further have dunite, norite, spinel troctonite, troctonite, and gabbro anorthosite (Burke 1986; Papike et al. 1991; Warren 2005; Fagan 2016; Heiken et al. 1991). The other regions, the dark areas, Mare, are composed of basalts that are rich in pyroxene (iron and magnesium). Olivine is an iron-bearing mineral, and ilmenite contains iron and titanium. Past the Apollo missions, it was found that the silicates (pyroxene, olivine, and feldspar) and oxides (including ilmenite, spinel) are the most common minerals in lunar rocks. Also, almost all the rocks fall into igneous category as they formed by cooling of lava. Some of the impact-generated fragments called breccia were also found, with the physical composition of broken and shock-altered fragments (clasts) and completely impact-melted material. Lunar soil is derived from rocks, remains of micrometeoroid bombardment and thermal, particulate, and radiation environments. It also leads to the formation of agglutinates (Papike et al. 1991; Harland 2008). From the absorption spectra measured through various missions to the moon, Ti3+, Fe2+, and Ti4+ ions are found in pigeonite, augite, olivine, and pyroxferroite. Absorption spectra can be a source for the characterization of oxidation states and transition metal ions. The presence of Cr3+ is also supported by few low oxidation states and reduction mechanisms mainly due to solar-wind

3

ion bombardment, meteorite impact, or magma genesis at depth (Bart 2007; Young 2017; Dhingra 2018). Telescopic, laboratory, and orbital measurements found that olivine causes the prominent absorption feature in reflectance spectra. Few intense absorption features were observed, which corresponds to high titanium contents. It also shows the molecular patterns. The Ti4+, Ti3+, Cr3+, Cr2+, Fe3+, and Fe2+ ions are the primary and detectable spectral features found in the nearinfrared-visible-near ultraviolet region as observed in electronic absorption spectroscopy for lunar surface (Papike et al. 1991). As the Ti-containing lunar soil matures, titanium got mixed into the lunar glass. It leads to intensive charge transfer, i.e., Fe2+, Ti4+, and the variation in the 300–600 nm wavelength range (Lucey et al. 1995, 2006). Due to a high amount of High-Ca pyroxene (> 50%), mare basalt has an absorption feature at 980–1000 nm. Depending on the variation in the content of olivine, the absorption band due to pyroxene will become wider, and the band center will shift toward the longer wavelength range. Absorption in the spectrum for lunar basalts varies highly with the content of olivine (0 to 20%)(McCord et al. 1981; King and Ridley 1987). Lunar highlands contain different proportions of feldspar minerals like anorthite. In the absence of magnesoferrous minerals, there will be no absorption feature in the spectra of anorthite. Also, due to plagioclase and low calcium pyroxene, the norite exhibits the characteristic absorption around 900–930 nm; however, the position of this band center varies with the amount of pyroxene (Williams and Jadwick 1980; Ulrich et al. 1981; Karr 2013). The rocks with minor HighCa whose absorption bands appear at 930–950 nm are due to norite, Gabbro containing feldspar, and High-Ca pyroxene located at 970–1000 nm. Dunite and troctonite, composed of olivine and feldspar, exhibit a wide multi-absorption feature near 1100 nm. KREEP has a composition similar to anorthosites, shows reflectance spectral characteristics closer to anorthosites (Warren and Wasson 1979; Warren 1989; Zou et al. 2004; Demidova et al. 2007).

A

4

Critical Absorption Features on the Moon Pyroxenes Absorption due to crystal filed transition in Ferrous ion (Fe2+) at M1 and M2 sites were observed at 478 nm, 1000 nm, and 2000 nm. However, a significant shift in maxima was also observed for this particular spectral feature. The feature near 505.05 nm shows the charge transfer processes for iron or titanium, ferrous, or ferric ions. Crystal field transitions in Cr3+ ions were noted at 450.50 and 600 nm. The significant footprints recorded by spectrometers are absorption bands for pyroxenes at 1000 and 2000 nm, due to ferrous ions. Band parameters like the relation between band center at 1000 and 2000 nm allow the detection of the ratio of iron and magnesium in pyroxene phases (Pieters et al. 2006; Klima et al. 2011). Olivine The absorption bands at 890–900 nm (M1), 1150 nm (M1), and 2050 nm are assigned to ferrous ions 1150 nm. Absorption at 720 nm and 1050 nm belongs to crystal field transitions in Cr2+ ions. A higher amount of olivine leads to a broader and shifted absorption band toward the long-wavelength range (Toksöz et al. 1974; Lin et al. 2020). Plagioclase Feldspar It was seen as a broad and polarized absorption band around 1250 nm and may be attributed to Fe2+ ions. The presence of Fe2+, Cr3+, Ti4+, and Ti3+ in lunar silicates is now well informed by the authors (Lucey et al. 1998, 2000). These conclusions were derived by examining the laboratory reflectance spectra of single crystals and whole-rock powders using diffuse reflectance methods. Critical absorption features are 300, 415, 750, 900, 950, and 1000 nm (Zou et al. 2004; Ohtake et al. 2009; Glotch et al. 2010; Isaacson et al. 2011).

Absorption Recorded from the Spectra of Apollo Samples Visible to Near-Infrared (300–2600 nm) For Apollo samples, researchers have used various grain sizes varying from a bulk sample,

Absorption, Absorption Spectrum in Lunar Studies

coarse-grained and grained (3 km to 200–300 m/s) can form secondary craters (Vickery 1986; Singer et al. 2014). Secondary craters are generally found in chains and clusters in the discontinuous and distal

Lunar Crater Ejecta

ejecta parts of the ejecta blanket, but isolated fragments can produce secondary craters that are essentially indistinguishable from primary impact craters. (Secondary craters are generally shallower than primaries, but over time they will begin to resemble degraded primary craters.) Ejecta Emplacement The ballistic flight trajectories of ejected fragments follow parabolic paths. When the ejected fragments strike the Moon’s surface, they possess the same velocity and impact at the same angle, and their momentum will scour the target surface and entrain target material within the newly formed ejecta blanket in a process known as ballistic sedimentation (Oberbeck 1975). The ballistic sedimentation process is important to consider, as it means that although the continuous ejecta blanket may appear as though it consists entirely of material excavated from the crater, a notinsignificant fraction of material in the ejecta blanket may be from local sources. The amount and severity of ballistic sedimentation depends on the size of the crater formed, ejection velocity and angle (and subsequent re-impact velocity and angle), and material properties of the substrate (Oberbeck 1975; Melosh 1989). For small craters ballistic sedimentation is less important than for larger, more energetic impacts (Melosh 1989). The material ejected by the impact process does not include material ejected from the full depth of the crater, but rather a horizon of material that is only one-third to one-half of the depth of the transient crater below the pre-impact surface (Fig. 1; Melosh 1989). Ejecta deposits around craters are heterogeneous, and the mixing process during excavation and ejection is incomplete (French 1998). The rocks found within the ejecta blanket represent the same assortment of rocks ejected, and preserve a recognizable sequence of the target rocks, albeit with the stratigraphy inverted. Near the crater rim, the sequence of inverted stratigraphy is often referred to as the overturned flap, a sharp recumbent fold with the hinge line marking the division between preexisting orientation and the inverted material. The preservation of an inverted stratigraphy over most of the radial length of the continuous ejecta

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blanket allows for the rudimentary reconstruction of original stratigraphic position and depth within the subsurface of the Moon (Shoemaker 1963). This principle was used by Apollo astronauts along sampling traverses of craters on the Moon. Material within the continuous ejecta blanket located farther from the rim is excavated from deeper within the crater, whereas near-rim material has been overturned from near the surface (Heiken et al. 1991; Shoemaker 1963; Shoemaker 1970). Additionally, the inferred depth of excavation is useful in remote-sensing studies where the composition of the ejected materials is not the same as the target surface. This is particularly true in cases of “dark-haloed craters” (e.g., Bell and Hawke 1984) where dark basaltic material was excavated by impacts though overlying bright, highlands materials overprinting mare (Fig. 2b, labeled dhc). Ejecta Thickness Ejecta blankets are thickest at the crater rim and gradually thin with radial distance. Not all of the material near the crater rim is ejecta, with a substantial portion of ~50% (Melosh 1989) to ~80% (Sharpton) of the volume of material coming from the stratigraphic uplift of the crater walls during the excavation process. The thickness of the ejecta blanket is somewhat poorly constrained and bounded by the volume of material from the preimpact surface to the crater floor. A rough approximation, based predominantly on explosion experiments, is given by equations defined by McGetchin et al. (1973), with a different interpretation that emphasizes the necessity of quantifying stratigraphic uplift given by Sharpton (2014). Ejecta Blanket Classification Lunar impact crater ejecta is typically classified by its distance from the final crater rim. References to crater ejecta are often divided into proximal and distal ejecta deposits and continuous and discontinuous ejecta blanket. Although often used interchangeably (i.e., proximal deposits and the continuous ejecta blanket referring to the same deposits), distinct designations exist for each of these regions. The proximal ejecta blanket is defined as the area within five crater radii (the

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Lunar Crater Ejecta

Lunar Crater Ejecta, Fig. 2 Copernicus crater (9.6 N, 20.1 E) viewed at (A) low-sun (LRO-WAC (Robinson et al. 2010) nearside morphology mosaic (643 nm)) and (B) high-sun (LRO-WAC normalized reflectance (albedo) map (643 nm)). Gray dashed line denotes the approximate extent of the continuous ejecta blanket. Herringbone shaped hills (hb) are often used to demarcate the transition from continuous to discontinuous ejecta facies. Secondary craters (sc) are typically found in chains and clusters roughly radial to the parent crater. Dark-haloed craters (dhc) can be seen in the albedo map and represent impacts that postdate the formation of Copernicus and have deposited dark material on top of the bright ray system

radius from the center of the crater to the final crater rim; French 1998) and encompasses the entire continuous ejecta blanket and the onset of the discontinuous ejecta blanket. The proximal ejecta contains most of the material (~90 vol%) ejected from the crater (Melosh 1989). The remaining ~10 vol% of material makes up the distal ejecta (>5 crater radii), which can travel tens to thousands of kilometers from the impact site (French 1998). Proximal deposits can be tens to hundreds of meters thick, depending on the crater size, although distal deposits are only a few centimeters to millimeters in thickness (French 1998). The continuous ejecta blanket is a region of thick ejecta (comprised of ~50 vol% of the total excavated volume of material) found within ~1 crater radius of the final crater rim (Melosh 1989;

French 1998). Morphologically, continuous ejecta blankets are the rough, blocky terrain of (generally) high albedo that contain radially emplaced grooves and scour marks and features such as boulder fields and impact melt flows, ponds, and veneer. The morphologic boundary between continuous and discontinuous ejecta occurs as a transition zone of mounded and hummocky terrain, with herringbone-shaped hills that point back to the crater center (Oberbeck and Morrison 1973) (Fig. 2a, labeled hb). Beyond this region the distal ejecta blanket begins at the onset of clearly defined secondary crater chains and clusters that occur as discrete radial features with interstitial regions having received little or no deposited ejecta material (hence its discontinuous nature). The discontinuous ejecta blanket and distal ejecta deposits mostly consist of secondary

Lunar Crater Ejecta

crater forming fragments and ray-forming particulates. Discontinuous and distal deposits do not have a well-defined outer boundary. The use of the terms proximal and distal deposits is most often (but not exclusively) used for descriptions of impactites (the general term for all rocks produced during an impact event (i.e., breccias, melt rocks, shock-metamorphosed target rocks)) (French 1998). The use of the terms continuous and discontinuous ejecta is most often used in morphology and morphometry studies of the ejecta blanket.

Morphology of Ejecta Facies Factors Affecting Ejecta Blanket Morphology Several factors can affect the final morphology of ejecta deposits, such as the impact angle of the projectile, the impact velocity, the material properties of the target (e.g., cohesiveness, composition, volatile content), preexisting topography, and the size of the impact crater formed (Melosh 1989; French 1998 (and references therein); Housen and Holsapple 2011). The impact angle affects the pattern and distribution of ballistic ejecta, with more oblique impacts concentrating ejecta downrange at impact angles >60 (from vertical) and the development of “forbidden zones” where ejecta materials do not land. These forbidden zones (Fig. 3) form up-range of the impact direction at impact angles 250 m/s 20% 44% 29% % Ejecta velocity >450 m/s 0% 22% 10% % Ejecta velocity >1800 m/s 0% 4% 1%

Measured actual 25–30 30 – – >1 1850 1850 ~20 ~40 –

~4 ~9 – – – – –

Notes: Assumes an impact density of 0.03 g/cm3, target (regolith) density of 1.5 g/cm3, 70 impact angle, 2.5 km/s Assumes impact temperature T ¼ 1200 K

a

Impact Flash A brief “flash-like” emission is detected in the LCROSS NIR spectrometer at 0.3 s after the calculated time of the impact with a 0.4 s rise time. It is also observed by the thermal camera (Schultz et al. 2010). This has been interpreted as a hot (1000 K) vapor cloud that emerged from the crater, corroborated by modeling of H2 gas observed by LAMP (Gladstone et al. 2010; Hurley et al. 2012). No flash is seen at visible wavelengths from LCROSS. Although an impact of a

2300 kg projectile at 2.5 km/s carries 7.2e9 J of kinetic energy, a large portion of this energy has been proposed to have gone into compressing the Centaur, driving material downward into a high porous (>70% porosity) material, and vaporizing volatiles, all of which acted to suppress the optical flash (Schultz et al. 2010; Hermalyn et al. 2012). Complex Plume A rise in the overall brightness in the LCROSS UV-visible and NIR spectrometers is seen and

Lunar Crater Observation and Sensing Satellite (LCROSS)

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Lunar Crater Observation and Sensing Satellite (LCROSS), Fig. 4 Key LCROSS observations of the plume created from the Centaur impact on October 9, 2009 in the Cabeus crater at the lunar south pole. (a) Ejecta cloud radiance, integrated over the bandwidth of the UV-visible and NIR LCROSS spectrometers, as a function of time after the Centaur impact. (From Colaprete et al. 2010, Fig. 1). (b) Spectral radiance difference from a baseline prespectra obtained with the LCROSS

UV-visible spectrometer indicate several emission lines. Time after impact is indicated in the legend. (From Schultz et al. 2010, Fig. 1D). (c) The expanding cloud as seen by the LCROSS visible imager is filled in and not hollow as per preimpact predictions. (From Schultz et al. 2010, Fig. 2A). (d) The resulting crater is observed by the LCROSS NIR imager. (Adapted from Schultz et al. 2010, Fig. 3A)

defines the arrival of the curtain into the sunlight a few seconds (1.1–3.1 s) after impact (Colaprete et al. 2010). Thermal fitting to the LCROSS spectra derives a hot (~300–1000 K) ejecta, consistent with LRO Diviner measurements made 90 s later (Hayne et al. 2010). The LCROSS UV-visible light curve, integrated radiance over 263–650 nm, peaks in brightness 18 s after impact (Fig. 4a), then decreases, remains flat for 150 s, and begins to rise ~180 s after impact (Colaprete et al. 2010). The spectrum also gradually gets bluer indicating the particles are becoming smaller, likely due to sublimination

and perhaps the breakup of loosely consolidated grains held together by volatiles (Heldmann et al. 2015). The cloud’s radiance is detected in the UV-visible and NIR spectrometers for the entire time the S-S/C acquired data, a 4-min, 16-s period (Fig. 4a). The LCROSS NIR radiance is observed to be 1/5th that of the UV-visible radiance, suggesting the cloud is not efficiently scattering NIR and/or there is higher absorption by species in the near infrared (Colaprete et al. 2010). A number of emission lines (CN, NH, NH2, CO2+, CS) are detected in the first 0.8 s by the

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Lunar Crater Observation and Sensing Satellite (LCROSS)

Lunar Crater Observation and Sensing Satellite (LCROSS), Fig. 5 As the LCROSS S-S/C and LRO instruments could only provide nadir imagery of the ejecta plume, observations with a more oblique view (from Earth) helps to constrain the ejecta morphology and dynamics. Shown is the dusty curtain, as surface brightness contour

maps, as viewed from the 3.5 m Apache Point Observatory on Earth at 10, 18 and 28 s after the Centaur impact. (From Strycker et al. 2013, Fig. 4). Strycker et al. (2013) has reprocessed data from previously reported Chanover et al. (2011) null result using improved removal of temporal variations

LCROSS UV-visible spectrometer (Schultz et al. 2010). These lines persist or are strengthened as greater amounts of ejecta reached sunlight (Fig. 4b). Their source is attributed as coming from the lunar regolith versus from any residual propellant in the Centaur. Sodium (Na) emissions at 598 nm and a line pair at 328 and 338 nm, attributed to silver (Ag), emerge along with H2S and H2O+. Sodium emission is also seen in the first 9 min with the Mc-Math Pierce Telescope in Arizona (Killen et al. 2010). LCROSS also detects hydroxyl (OH) in emission at 309 nm, attributed to the breakdown of H2O by thermal dissociation or excited OH desorbed from grain surfaces. OH emission shows increase in the first 50 s, decays from 50 to 90 s, rises again 90 s, and remains elevated out to ~230 s (Colaprete et al. 2010). Images from the LCROSS visible camera show the ejecta lasting tens of seconds and expanding from ~4 km (at 8 s) to ~8 km (at 20 s) in diameter before disappearing from view about 42 s when it dropped below the camera’s sensitivity threshold. The two on-board NIR cameras also capture images of the expanding cloud, but for less time, as the cloud brightness is much fainter

for their dynamic ranges (Schultz et al. 2010). From the nadir point of view (Fig. 4c), the curtain is observed to be relatively symmetrical and does not show a central “hole” as predicted by laboratory tests and computer modeling by Schultz 2006 and Korycansky et al. 2009. The dusty curtain (Fig. 5) is also seen at >4 km above crater rim (~12 km above crater floor) by the 3.5 m Apache Point Observatory, New Mexico (Strycker et al. 2013). A broad continuum from 130 to 170 nm and strong emission at 185 nm are seen by LRO’s LAMP ultraviolet spectrometer in the 30–60 s after the impact. A cloud of molecular H2 and CO, with Hg, Ca, and Mg, expanding from an initial temperature of 1000 K fits the observed UV spectra. LAMP’s slit was aligned above the lunar limb and able to observe ~27–38 km above the LCROSS impact site from ~103 km away. From their viewpoint, the emission is brightest at ~45 s after the Centaur impact, implying a bulk velocity of ~1.9 km/s at least, for the species seen. LAMP observes emission earlier indicating the plume has a range of speeds, with fastest ~4.6 km/s (Gladstone et al. 2010).

Lunar Crater Observation and Sensing Satellite (LCROSS)

The observed H2 is consistent with the formation of H2 from adsorbed H present in the local regolith rather than from the H2 created from the photolysis of water (Gladstone et al. 2010). It is also much higher than the hydrogen estimated by the LRO neutron spectrometer (Mitrofanov et al. 2010). This might imply that hydrogen is concentrated in areas smaller than the 10 km footprint of those neutron measurements. Alternatively the H2 signatures seen by LAMP could have come from a much greater depth (Hurley et al. 2012). Table 3 summarizes the relative concentrates of volatile species observed by LCROSS (Colaprete et al. 2010) and LAMP (Gladstone et al. 2010). LCROSS certainly hit a very rich and diverse cold trap. H2 and OH (g) may have been products driven by heat created during the impact (Metzger et al. 2020). Per Poondla et al. (2020), OH(g) could have also been produced from water vapor (from sublimating ice grains). All other volatiles could have been in their solid form at the basin of the PSR. Their origin could be from comets, interstellar medium, or be volcanic (intrinsic to the Moon). Volatile origin, entrapment, mobility, retention, and escape remain an active area of research in planetary science. Postimpact Plume Modeling

The LCROSS impact has liberated many species from the target site Cabeus. However, the plume evolution appears to behave differently depending on the species been examined. This has led to new work in complex plume modeling by several research groups. Most prior laboratory studies with vertical impacts into lunar regolith-like (sand) surfaces have demonstrated a narrow range (45 –55 ) of curtain ejection angles (see Cintala et al. 1999; Anderson et al. 2003; Hermalyn and Schultz 2010). None of these models could fit the LCROSS observed plume at the visible wavelengths. Postimpact laboratory experiments would reveal that hollow projectiles into the same target present a very different scenario characterized by two plumes: (1) a high-angle component emerging initially at high speeds (>1 km/s) and decreasing in speed with time and (2) an early-time appearing and long-lived low-angle ejecta

515 Lunar Crater Observation and Sensing Satellite (LCROSS), Table 3 Volatiles in the LCROSS ejecta. (From Colaprete et al. 2010; Gladstone et al. 2010) Compound Water Hydrogen sulfide Sulfur dioxide Ammonia Carbon dioxide Ethylene Methanol Methane Hydroxyl Carbon monoxide Calcium Hydrogen gas Mercury Magnesium

Symbol Concentration (% by weight) 5.5 H2O H2S 1.73 SiO2 NH3 CO2 C2H4 CH3OH CH4 OH CO

0.61 0.32 0.29 0.27 0.15 0.03 0.0017 0.000003

Ca H2 Hg Mg

0.0000008 0.0000007 0.0000006 0.0000002

component. This new two-plume model fits the LCROSS data better where the low-angle plume originates from 85 vol. % An) exhibiting a plutonic or relict texture (Dowty et al. 1974a, b). FAN rocks are the most common pristine highland rock, constituting nearly 50% of the outer highland crust. Compared to terrestrial rocks having high Ca/Na ratio and other pristine highland rocks, olivine and pyroxene in FAN are Fe-rich, hence called as Ferroan (Hiesinger and head 2006; Lucey et al. 2006 and references therein). FAN rocks are believed to be the remnants of the original floatation crust of the Moon formed from the magma ocean around ~4.6–4.4 Ga ago. FAN rocks are commonly found as monomict breccia. Plutonic remnants are seldom reported from Moon. High plagioclase content of the rocks indicates their formation by separation of early formed plagioclase crystals

Lunar Crust, Chemical Composition

from the residual melt. The second most abundant mineral is pyroxene, commonly low-Ca pigeonite which in some cases, exsolved to form orthopyroxene host crystal and lamellae of augite (Warren and Wasson 1978). Augitic pyroxenes were also seen in some samples (Ryder 1985). Primary igneous origin of the augitic pyroxene was inferred as it was found in cumulate textured ferroan gabbro (Warren et al. 1983d). FAN rocks also consist of olivine, ilmenite, Fe-metal (kamacite), troilite, silica phase, and Cr-Fe spinel. Ferroan anorthosites are estimated to be 4.56–4.29 Ga old. Sm-Nd dates gave 4.44  0.02 Ga age for FAN (Lugmair 1987) while Pb-Pb dates revealed an age of 4.51  0.01 (Hanan and Tilton 1987). Recent orbital remote sensing based studies have detected pure anorthosites (PAN) with plagioclase abundance nearly 100 vol. %, mainly on the central peaks of impact craters (Ohtake et al. 2009). It was also shown that the pure anorthosites are broadly distributed on the nearside and farside of the Moon (Donaldson Hanna et al. 2014). These pure anorthosites are found in the thicker regions of the crust, mainly in the surrounding regions of impact basins. Global distribution of pure anorthosites combined with the GRAIL (Gravity Recovery and Interior Laboratory) data suggested a thickness of at least 30 km for the primary anorthositic crust of the Moon (Donaldson Hanna et al. 2014). Mg-rich rocks: Mg-rich rock types on Moon show range of compositions from olivine-rich dunites to pyroxene and Na-rich plagioclase rocks (sodic ferrogabbro). Mg-rich rocks are generally coarse-grained highland igneous rocks which can be distinguished from the FAN in terms of their high Mg/(Mg + Fe) and Na/(Na + Ca) ratios (Warner et al. 1976). Plagioclase, pyroxene, and olivine constitute 99% of the Mg-rich rocks. Plagioclase is more Na-rich and mafic silicates are more of magnesian and abundant. Gabbros, norites, and troctolites are the common Mg-rich rock types on Moon. Plagioclase and pyroxene are the major minerals in gabbro (clinopyroxene+plag) and norites (orthopyroxene+plag) while troctolites are primarily made up of plagioclase and olivine. Trace minerals in Mg-rich rocks include Cr-Fe spinel, Fe-Ni metal

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(kamacite, and taenite), whitlockite, apatite, baddeleyite (ZrO2), K-Ba feldspar, and pyrochlore ([(Na, Ca)2(Nb-Ta)2O6(OH, F)]) (Dymek et al. 1975). Mg-Al spinel, ilmenite, troilite, farringtonite (Mg3(PO4)2), armalcolite, zircon, zirconolite, and silica minerals are the commonly found accessory minerals in Mg-rich rocks (Prinz and Keil 1977; Ryder and Norman 1979; James 1980). Most of the Mg-rich rocks fall in three groups: (a) highly magnesian anorthositic troctolites, (b) moderately magnesian norites, (c) gabbro norites containing typical low-calcium pyroxene and/or much higher high-calcium pyroxene. Most Mg-rich rocks are brecciated; however, igneous textures are found to be preserved in relatively undeformed rock fragments indicating their origin as igneous cumulates. Mg-rich rocks would have been formed likely after the solidification of the primordial magma ocean. Coarse granular texture in few rock samples suggested thermal metamorphism for a prolonged period or slow cooling and equilibration (Gooley et al. 1974; Dymek et al. 1975). Diversity in the bulk composition of Mg-rich rocks indicates the removal or accumulation of different minerals during fractional crystallization. Troctolites are commonly accumulated from the melt evolving along plagioclase-olivine boundary whereas norites accumulate probably from the same melt at later stages or from a more silica-rich melts evolving along the plagioclase-pyroxene boundary. Mg-suite rocks are about 4.46–4.18 Ga old coinciding with the estimated age of FAN-suite rocks but appear to have somewhat younger ages (Shearer et al. 2006). Sm-Nd dates on troctolites 76535 gave an age of 4.26 Ga whereas Rb-Sr dates showed a much older age of 4.51 Ga. The formation age of these rocks must be greater than 4.2 Ga despite the discrepant ages provided by radiometric dating methods. The range of radiometric ages indicates that earlier Mg-suite magmatism would have occurred simultaneously with some of the early formed Ferroan anorthosites (cf. Heiken et al. 1991). However, the near-contemporaneous crystallization of at least some of the Mg-suite rocks and Ferroan anorthosites contradicts the basic concept of magma ocean model which involves the formation of Ferroan anorthosites followed by the intrusive magmatic activities

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Lunar Crust, Chemical Composition, Fig. 4 (a) Image of the Apollo 16 rake sample, 65785, with a pink spinel troctolite clast in it (labeled as PST) (Credits: NASA

Lunar Crust, Chemical Composition

S72–48821). Scale bar is in mm. (b) Thin-section image of the lunar meteorite ALHA81005 showing the pink spinel (pink-grains) in polarized light (From Gross and

Lunar Crust, Chemical Composition

producing plutonic Mg-suite rocks at later stages. Considering the overlapping ages of these rock types on Moon, an alternative model to the global magma ocean concept was favored involving serial magmatism to be a probable mechanism in the crust-building process (Walker 1983; Longhi and Ashwal 1985; Longhi 2003; Shearer et al. 2006; Borg et al. 2011). However, the exposures of pure crystalline anorthosites in the impact excavated regions (Ohtake et al. 2009; Donaldson Hanna et al. 2014) are consistent with the formation of the primary crust by floatation of plagioclase in a cooling magma ocean. Hence, it was inferred that the complexity of crustal rocks now found on Moon is a consequence of the solidification of the magma ocean and redistribution of the primary crustal materials due to violent impacts and other possible events like localized mantle overturn which brought up the Mg-suite intrusives and concentrated in the Procellarum KREEP terrane or at the base of the crust (Jaumann et al. 2012). Apart from these Mg-rich rocks, a new rock type, pink spinel anorthosites (PSA), was also discovered on the Moon recently using Chandrayaan-1 data (Pieters et al. 2011; Pieters et al. 2014). PSA are primarily composed of Mg-Al spinel, plagioclase, and < 5% mafic minerals such as olivine and pyroxene. PSA are found to be associated with central peaks and rims of impact craters and inner ring of the basins (Fig. 4c). Previous detections of pink spinel (pink-hued; Magnesium Aluminum Oxide) were reported from Apollo and Luna samples (Fig. 4a, b). Early detections of pink spinels were commonly found in troctolites containing plagioclase and olivinepyroxene (>8 vol. %). However, recently discovered PSA contains no or lower percent of mafic components (90 % plagioclase, and had textures indicating an origin as a cumulate rock rather than fragments from the basaltic lavas also found at the landing site (Smith et al. 1970a; Wood et al. 1970a). Secondly, the robotic Surveyor 7 lander had measured a soil with similar composition to that found at the Apollo 11 site (Patterson et al. 1969; Phinney et al. 1969; Wood et al. 1970a), implying this abundance of anorthosite was widespread across the lunar surface. Thirdly, both Wood et al. (1970a) and Smith et al. (1970a) recognized a concentration mechanism was needed to accumulate anorthosite in the lunar crust. Both groups suggested this was achieved by fractionation and

Lunar Magma Ocean Theory, Origins, and Rationale

flotation of anorthosite on a denser Fe-rich magma layer, which by necessity must have covered most of or the entire Moon. Lastly, Wood et al. (1970a) argued based on early gravity field data from the Lunar Orbiter missions (O’Keefe 1968), which suggested a 25 km thick lunar crust, that if the crust was predominantly anorthosite derived from fractional crystallization of a basaltic magma, the amount of anorthosite this implied in the crust required melting of a substantial fraction of the Moon. The fact that two groups independently proposed a model for the differentiation of the Moon, within 6 months of the return of a mere 22 kg of material, that is still accepted at present day is remarkable. The LMO model proposed by the Harvard and Chicago groups has certainly been refined and improved in the years following Apollo 11 as more samples, data, and improved analytical techniques became available, but the basic framework has endured. Below is a summary of the three primary lines of evidence that argue for large-scale melting of the Moon and a brief description of the reasoning that underlies them. Other lunar datasets not detailed here have been argued to support the LMO hypothesis, and there have been challenges to LMO theory that have arisen based on seemingly incongruent data (see Shearer et al. 2006 for a discussion of complexities and challenges to LMO theory). However, the LMO hypothesis remains broadly accepted in the planetary science community due in large part to the inability of other models to explain the observations discussed in the following sections.

Geochemical Evidence for Large-Scale Melting of the Moon A Global, Ancient, Plagioclase-Rich Crust The first strong line of evidence that the Moon underwent a global, high-temperature differentiation event is the composition of the lunar highlands crust. Wood et al. (1970a) first examined the soil samples returned by Apollo 11 and the rock fragments they contained. They noted that of the 1676 fragments examined, 61 were fragments of

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anorthosite, which was a finding that they did not expect. Although the anorthosite fragments ranged in purity from true anorthosites (>90 % plagioclase) to anorthositic gabbros ( 500 km, primordial lunar mantle is present (Image is modified from Shearer and Papike (1999))

Lunar Magma Ocean, Size

Conclusions To date, differentiation and crystallization of a magma ocean on an early-formed Moon are arguably the most elegant (Shearer and Papike 1999) framework in which current lunar lithological, mineralogical, geochemical, geochronological, and geophysical data should be evaluated. Constraints on the extent of melting associated with a primordial LMO on a young planetary body have however, not been consistent, ranging from 90% olivine), troctolites (olivine and plagioclase), norites (plagioclase and orthopyroxene), and gabbronorites (plagioclase and pyroxene). These rocks represent later intrusions into the highlands crust of ferroan anorthosite at 4.43-4.17 Ga (Taylor et al. 1993), which are slightly younger than the ferroan anorthosites (4.5-4.3 Ga Hiesinger and Head 2006).The alkali suite contains alkali anorthosites with relatively sodic plagioclase (An70–85), norites, and gabbronorites with more iron in mafic minerals than the magnesian suite. The isotopic age for alkali suite is similar to the magnesian suite.

Lunar Meteorites

Comparing to highland rocks, mare basalts are enriched in FeO and TiO2, depleted in Al2O3. Mineralogically, mare basalts are enriched in olivine and pyroxene, especially clinopyroxene, and depleted in plagioclase. Among the classification schemes, TiO2 content is the most useful discriminator to classify lunar mare basalts (e.g., Neal and Taylor 1992; Papike et al. 1998). Using TiO2 concentration, mare basalts could be classified into three groups: high-Ti basalts (>9 wt%), low-Ti basalts (1.5–9 wt%), and very low-Ti (VLT) basalts (85.5 lat targeted >10% of lunar surface

Coverage of data sets Global coverage; Polar grid 0.001 latitude and 0.04 longitude; High accuracy 50 m of global geodetic grid

Lunar Reconnaissance Orbiter (LRO) Mission, Table 2 Details and specifications of instruments. (Tooley et al. 2010; Smith et al. 2010a; Spence et al. 2010; Mitrofanov et al. 2010; Paige et al. 2010; Nozette et al. 2010; Gladstone et al. 2010; Litvak et al. 2012a, b)

Lunar Reconnaissance Orbiter (LRO) Mission 781

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(1) maps the entire lunar surface in the far ultraviolet; (2) searches for surface ice and frost in the PSRs; (3) feasibility of using starlight and sky glow for future surface mission

LAMP

Science goals (1) global high-resolution hydrogen distribution maps; (2) search for evidence of water ice; (3) provides information about the lunar radiation environment (thermal, epithermal, and high-energy neutron fluxes)

(1) day/night thermal mapping measurements of entire lunar surface; (2) identify cold traps and potential ice deposits; (3) determine rock abundances globally

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Mass and power (avg) 25.8 kg, 11.6 w

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Instrument LEND

Lunar Reconnaissance Orbiter (LRO) Mission, Table 2 (continued)

Far UV: 52~187 nm, 3.5 nm sampling; 260 m spatial resolution (footprint)

9 channels (mm): 0.35–2.8 (high sensitive), 0.35–2.8 (low sensitive), 7.55–8.05, 8.10–8.40, 8.38–8.68, 13–23, 25–41, 50–100, 100–400 20~400 K, resolution of 5 K; Spatial resolution 400 m

Parameters (at the altitude of 50 km for all instruments) Nine neutron sensors, four bands: Thermal neutron sensors (STN1– STN3) 5 years; these produced a power source of >70 Welectric. In an attempt to turn the Apollo seismometers back on in 1986, it was found that there was not enough operating power. The seismic signals recorded by the PSEs were very different from those seen on Earth in a number of characteristics, such as duration, onset, and shape of the envelope. For example, the signals from the impacting lunar modules lasted much longer than would have been the case on Earth. The Apollo PSEs defined four categories of natural seismic events: shallow moonquakes, deep moonquakes, meteorite impacts, and thermal moonquakes. Each of these moonquakes produces distinctive seismograms. Latham et al. (1972) reported that deep moonquakes (~800 km) repeated in monthly cycles triggered by lunar tides. It appears that such deep events originate from distinct regions within the lunar mantle; more than 3000 deep moonquakes have been assigned to 109 separate hypocentral regions and more recent work has increased this number. In addition to the repeated moonquakes, moonquakes swarms also occur, maybe as frequently as one every 2 h over intervals lasting several days. A swarm was defined as 8–12 seismic events per day compared to the usual 1-2 events per day.

The Sources of Seismicity Over the 8 years of Apollo passive seismic recording, the largest recorded moonquakes have Earthequivalent magnitudes of about 4 (Neal et al. 2004). The sources of seismicity on the Moon include the follow four categories:

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Deep Moonquakes The most abundant type with >7,000 events were recognized originating from 700 to 1,200 km depth. These small-magnitude events (5 magnitude. Exact focal depths are unknown because all recorded events were outside the limited network. Indirect evidence suggests depths between 50 and 200 km. They are not correlated with tides but may be associated with boundaries between dissimilar surface features. Although there were only about four or five shallow moonquakes per year during the period of lunar seismic monitoring, they account for most of the seismic energy released in the Moon (Nakamura 1980). Meteoroid Impacts While most of the energy of an impact is expended excavating a crater, some is converted to seismic energy. Between 1969 and 1977, >1,700 events representing meteoroid masses of 0.1–100 kg were recorded. Events generated by smaller impacts were too numerous to be counted. Seismic events due to meteoroid impacts vary widely in energy. Meteoroid impacts of all

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energies tend to be most common when meteoroid showers peak (Dorman et al. 1978), particularly among the largest meteoroid impacts that tend to occur in the months of April through July. The largest recorded impacts, in July 1972 and May 1975, represented meteoroids of about 5 t. In all, seven meteoroid impacts of 1 t or more were observed within 5 years during lunar seismic monitoring (Latham et al. 1978; Dorman et al. 1978).

The Application of the Seismic Data As a result of the Apollo Passive Seismic experiments, the release of seismic energy from the Moon is commonly assumed to be small, only ~2  1010 J/year compared to Earth’s 1017 to 1018 J/year (Goins et al. 1980). However, larger but rarer moonquakes may not have occurred during the 8 years of lunar seismic monitoring, and the actual average lunar seismic energy could be as high as 1014 J/year (Nakamura 1980). In either case, the Moon’s low seismic activity, coupled with extremely low elastic wave propagation losses (e.g., low attenuation, sometimes referred to as “high Q”), mean the Moon is an extremely quiet place, even though seismic events and the resulting elastic “sounds” carry for long distances through the rock and soil with great clarity. (Note that there is no sound transmission through the air as the atmosphere is too thin.) The Apollo 11 passive seismic instrument clearly recorded Astronaut Armstrong climbing the ladder into the LM. This sensitivity to seismic energy because of low attenuation gave rise to the phrase “the Moon rings like a bell,” as seen in the characteristically long seismic signatures of moonquakes and of meteoroid impacts on the Moon.

Future Work Need for a Global Lunar Seismic Network This is required to locate the origins of the different types of moonquakes, especially those that could compromise a Moon base. (1) A statistical analysis of meteorite impact sites is required to

Lunar Surface, Seismic Properties

determine if the Moon base site has a statistically low probability of receiving a sizeable meteoroid impact. (2) Understanding the nature and location of shallow moonquakes is required so the Moon base site is not in a seismically active area. These examples are prudent in terms of safety and to protect the required investment. Required Technological Advances An international group of scientists has been investigating the challenges of establishing a global Lunar Seismic Network. A modest network requires 8 seismometers to be deployed around the Moon and be active for 5–7 years. Soft and hard landing options have been explored. Both have their limitations, which require technological advances in 3 inter-related areas: (1) Deployment – mass must be reduced through hardware miniaturization; (2) Hardware – needs to be more robust such that the mass required for deployment can be reduced; (3) Powerdevelopment of robust mini radionuclear thermoelectric generators (RTGs) that can maintain a power supply of 3–5 watts over 5–7 years yields a huge mass reduction. Developing such technology for a LuSeN-type mission will allow for similar exploration of Mars and beyond.

Cross-References ▶ Internal Structure/Mantle Motions of the Moon ▶ Regolith Physical Properties

References Dorman J, Evans S, Nakamura Y et al (1978) On the timevarying properties of the lunar seismic meteoroid population. In: Proceedings of the nineth lunar planetary science conference, pp 3615–3626 Goins NR, Dainty AM, Toksoz MN (1980) Seismic energy release from the Moon (abstract). Lunar Planet Sci XI:336–338 Latham GV, Dorman HJ, Horvath P et al (1978) Passive seismic experiment: a summary of current status. In: Proceedings of Lunar and Planetary Science Conference, pp 3609–3613 Latham GV, Ewing M, Dorman J et al (1972) Moonquakes and lunar tectonism. The Moon 4:3–12

Lunar Tectonics Nakamura Y (1980) Shallow moonquakes: how they compare with earthquakes. In: Proceedings of the eleventh lunar planetary science conference, pp 1847–1853 Neal CR (2005) The Importance of Establishing a Global Lunar Seismic Network. Space Resources Roundtable VII: LEAG Conference on Lunar Exploration. LPI Contribution No. 1287, pp 70 Neal CR, Banerdt WB, Chenet H et al (2004) The Lunar Seismic Network: Mission Update. In: Proceedings of Lunar and Planetary Science Conference, pp 2093– 2094

Lunar Tectonics Regular Wave Tectonics of the Moon G. G. Kochemasov IGEM of the Russian Academy of Sciences, Moscow, Russian Federation

Despite numerous random features on the lunar surface tied to irregular impact processes, there is evidence of regular wave-borne structures of various scales. The most important among them is the largest, covered with dark basalts, depression known as the Procellarum Basin. It is a characteristic feature of the lunar nearside clearly visible during the Full Moon (Fig. 1). Recently, this large feature was considered as a typical impact. Systematic gravity surveys, however, established that the Procellarum Basin is a tectonic formation of deep origin (Andrews-Hanna et al. 2014). The antipodal farside is composed mainly of “continental” highland formations. The dichotomous opposition of the highland and lowland hemispheres, which is typically observed in the structures of planetary bodies of various classes, sizes, and compositions, is a characteristic phenomenon of wave origin. The most fundamental wave mode divides a globe into two hemispheres: uplifted and subsided ones. The Moon is no exception. A natural consequence of this division is in different compositions of the antipodal segments-hemispheres aimed to level the angular momentum of tectonic blocks with different planetary radii. Angular momentum of the

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subsided nearside is supported by the dense basalt fill; enhanced angular momentum of the uplifted farside is diminished by the less dense anorthosites. Ubiquitous dichotomous tectonics of various planetary bodies exhibits the wave nature of their structuring. The fundamental wave-1long 2πR has harmonics creating their own tectonic blocks. Thus appear tectonic sectors (creation of the wave-2 long πR) superposed on the dichotomy. On the Moon very striking South Pole-Aitken Basin of the far hemisphere is a representative of tectonic sectors though many scientists think it is a huge random impact (Fig. 1). Along with its deepest lunar hypsometric marks (more than 8 km), it coincides with the deepest lunar geoid minimum (Zuber et al. 2013). It is remarkable that this lunar geoid minimum completely tectonically coincides with the terrestrial deepest geoid minimum of the Indian Ocean. Both minima and hypsometric “holes” are covered with dense basalts and bordered at the north by the highest continental blocks (Fig. 2). This layout strengthens tectonic similarity between two coupled cosmic bodies (planet-satellite) and raises fundamental natural science question of energy source for planetary structuring. Between the two blocks of primary tectonic importance (Procellarum Basin and SPA Basin – 2πR- and πR-structures) occurs a remarkable “ring” of the Mare Orientale (πR/2-structure). Its tectonic terrestrial analogy is the Malay Archipelago “ring.” The global sequences of three tectonic blocks of various sizes on the Moon and Earth emphasize tectonic regularity of cosmic bodies organized by the planetary wave process (Kochemasov 2017). Waves warping planetary bodies originate because the bodies move in Keplerian elliptical orbits with periodically changing accelerations. Arising forces (bodies’ mass multiplied by acceleration ¼ force) applied to a cosmic body warp them in various harmonic waves spreading in four ortho- and diagonal directions. Produced by crossing waves, regularly disposed rising and falling tectonic blocks make a regular tectonic pattern. Segments-hemispheres (2πR-structure), net of sectors (πR-structure), chains, and

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Lunar Tectonics

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60’ 30’

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Lunar Tectonics, Fig. 1 Tectonic triads. Comparison of Earth’s geography and Moon’s hypsometry. On the right, schematic sizes and relative dispositions of terrestrial (above) and lunar wave-borne tectonic features. Pacific Ocean and Procellarum Ocean on the right – 2πRstructures. Indian Ocean and SPA Basin on the left – πR-

Lunar Tectonics, Fig. 2 Moon’s (to the left) and Earth’s geoids (Topographic Map of Earth’s Moon. Credit: from the Japanese Kaguya mission. Wikipedia.qwika. com)

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structures. Malay Archipelago and Mare Orientale at the center – πR/2-structures. Equator (axis of rotation) positions at present (Earth) and in ancient Moon (Topography of the Moon from the Chang’E-1 laser altimeter data CAS Oceanearth)

Lunar Tectonics

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Lunar Tectonics, Fig. 3 Gravity map of the Moon. LPOD-Sept6-09 (www2.lpod.com)

L grids of tectonic granules (πR/2-structures) superimposed on these larger blocks (2πR and πR) are elements of the regular wave tectonics (Kochemasov 2017). Definite and very impressive chains of negative forms, Maria, intersect both lunar hemispheres (Fig. 3). At present they are inclined, but in the past (when the axis of rotation was ~30 from the present (Garrick-Bethell et al. 2014), they made one belt along the equator. This alignment mimics the equatorial belt of eight terrestrial rings – one more reason to see the similarity of the Moon’s and Earth’s tectonics. The chain crossing the far highland hemisphere consists of the Mare Moscoviense, Freundlich-Sharonov, DirichletJackson, Hertzsprung, and Orientale. The chain on the nearside includes Mares Imbrium, Serenitatis, Crisium, and Smythii (Fig. 3). Roughly equal distances between Maria are very evident, which favorably indicate their wave origin. The nearside objects are marked by pronounced mascons due to nearness to surface of dense basaltic mantle.

Cross-References ▶ Internal Structure/Mantle Motions of the Moon ▶ Kaguya (SELENE) Mission ▶ Lunar Landscape, Highlands ▶ Lunar Landscape, Maria ▶ Lunar Magma Ocean ▶ Lunar Magma Ocean, Comparison to other Planetary Magma Oceans ▶ Lunar Mare Basalts, Stratigraphy of ▶ Lunar Surface, Gravity Field ▶ Lunar Tectonism, History of ▶ Lunar Terrane Tectonics ▶ Surface of the Moon, Distribution of Materials and Structures ▶ Topographic Studies of the Moon ▶ Volcanic Processes on the Moon

References Andrews-Hanna J, Besserer J, Head J III et al (2014) Structure and evolution of the lunar Procellarum region as revealed by GRAIL gravity data. Nature 514:68–71

842 Garrick-Bethell I, Perera V, Nimmo F, Zuber M (2014) The tidal-rotational shape of the moon and evidence for polar wander. Nature 512:181–184 Kochemasov G (2017) New planetology and geology: tectonic identity and principal difference of terrestrial oceans and lunar basins. New Concepts Global Tectonics (NCGT) J 5:131–133 Zuber M, Smith D, Watkins M et al (2013) Gravity field of the moon from the gravity recovery and interior laboratory (GRAIL) mission. Science 339:668–671

Lunar Tectonism, History of Paul K. Byrne Planetary Research Group, Department of Marine, Earth, and Atmospheric Sciences, North Carolina State University, Raleigh, NC, USA

Definition The temporal sequence by which tectonic landforms on the Moon were formed and what the types and spatial distributions of those landforms tell us of their formation mechanisms, with implications for the thermal evolution of the body.

Introduction The Moon shows no evidence for having experienced plate tectonics, such as divergent or convergent plate boundaries, volcanic arcs, subduction trenches, or extensive rift zones. Instead, the Moon experienced an early phase of differentiation during primordial cooling that segregated buoyant plagioclase from a global magma ocean and led to the development of a feldspathic crust (e.g., Wood et al. 1970; Green et al. 1971). This crust, likely mechanically uniform and less dense than the underlying mantle, rendered the Moon a one-plate or stagnant-lid planetary body (e.g., Solomon and Head 1979; Hiesinger and Head 2006). With a thick, rigid outer layer hundreds of kilometers thick (e.g., Nakamura et al. 1973; Wieczorek et al. 2006), no interior motion of the mantle could be transferred to the surface in the manner of plate tectonics on Earth.

Lunar Tectonism, History of

Major global-scale tectonic stresses on the Moon therefore originate not from convection but from conduction: the body has experienced some reduction in volume in response to the steady cooling of its interior to space (e.g., Solomon 1978). This process, termed global contraction, has been accompanied by tidal forces (e.g., Hiesinger and Head 2006), impact-induced tectonism (e.g., Johnson et al. 2016), basinloading effects (e.g., Solomon and Head 1979), and shallow magmatic intrusion (e.g., Schultz 1976). These mechanisms are inferred to have operated on the Moon because of the specific types and distributions of tectonic landforms recognized on the body, which primarily include graben, wrinkle ridges, and lobate scarps. Graben are linear topographic lows that consist of a down-dropped block bounded by two normal faults that dip toward one another (i.e., that are antithetic) (e.g., Golombek 1979). They are interpreted as reflecting crustal extension perpendicular to the direction in which they are oriented. In contrast, wrinkle ridges are positive-relief landforms: broad, steep-sided, but low-relief rises that are generally symmetric in cross section and which occur widely across the lunar maria (e.g., Bryan 1973). They are generally regarded as corresponding to folds underlain by shallowly dipping thrust faults (e.g., Mueller and Golombek 2004). The morphology of wrinkle ridges attests to crustal shortening, again directed perpendicular to the orientation of the ridges themselves. Finally, lobate scarps are another type of shortening structure, similar in size to lunar wrinkle ridges but much more asymmetric in cross section (e.g., Masursky et al. 1978), but again suggestive of folding atop shallow thrust faults (e.g., Watters and Johnson 2010). Crucially, it is the relationships between these landforms and with other units on the Moon, and with each other, that provide insight into the time frames over which these processes have operated. The law of superposition holds that the uppermost unit of a geological sequence is generally the youngest and vice versa. This approach has been refined for Earth over the last several centuries and is widely applied to planetary geology. Where a tectonic structure is observed to crosscut a surface unit, for example, it is probable that the structure

Lunar Tectonism, History of

formed after that surface was emplaced. Similarly, volcanic flows embaying a tectonic structure, say, suggest that the structure was in place before the volcanic activity. And where one structure appears to offset or otherwise affect another, such as a wrinkle ridge deforming the floor of a graben, then a period of crustal shortening likely followed a phase of extensional tectonism. Through collating such observations for the entire Moons, it is possible to deduce a first-order history of lunar tectonism.

Extensional Structures Lunar graben are concentrated almost wholly on the lunar nearside, with but a few structures on the farside, associated with the Orientale and Schrödinger basins (Scott et al. 1977; Shoemaker et al. 1994; Watters and Johnson 2010). Indeed, most graben are situated within or immediately outside large basins (crossing into the lunar highlands) and show either dominantly basin-radial or basin-concentric orientations. Their close spatial association with impact basins implies that many graben are the result of impact-induced stresses (e.g., Schultz and Gault 1975; Ahrens and Rubin 1993). Moreover, the emplacement of mare lavas within these basins induces peripheral extensional stresses, further promoting extensional structures particularly at basin edges (Solomon and Head 1979; Wilhelms 1987; McGovern and Litherland 2011). Observations of lavas infilling graben, such as those around the Humorum basin (Platz et al. 2015), indicate that the lavas there were emplaced after the crust was extended. In fact, most graben on the Moon appear to have formed prior to about 3.6 billion years ago (Lucchitta and Watkins 1978; Solomon and Head 1979), with little evidence for widespread crustal extension since then. Extensional structures, albeit local in scale, do occur elsewhere on the Moon. For example, numerous impact craters feature floors crossed by extensive systems of graben (Schultz 1976; Jozwiak et al. 2012), with these structures creating complex polygonal patterns. Formation via viscous relaxation (Hall et al. 1981) or magmatic

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intrusion (e.g., Wichman and Schultz 1995; Dombard and Gillis 2001) has been proposed to account for these landforms, with a general consensus leaning toward igneous activity over viscous deformation (e.g., Jozwiak et al. 2015). In any case, although these structures occur in craters that formed throughout much of lunar history, the close geographic correlation between floorfractured craters and the lunar maria suggests that this limited crustal extension tracks approximately with the span of mare volcanism (generally Lower Imbrian to late Eratosthenian: Jozwiak et al. 2015), ending about 1.5–1 billion years ago. Curiously, the Gravity Recovery and Interior Laboratory (GRAIL) mission provided evidence for a population of ancient, gigantic extensional structures buried deep within the lunar interior (Andrews-Hanna et al. 2013). Manifest as linear positive gravity anomalies hundreds of kilometers long and ~5–40 km wide, these structures are likely enormous dikes. They have little to no surface expression, however, and crosscutting relations indicate that they are pre-Nectarian to Nectarian in age (i.e., predating the end of the late heavy bombardment of the Moon) (Andrews-Hanna et al. 2013). The locations and geometries of these structures, then, imply an early phase of considerable crustal extension – corresponding to an increase in body radius of between 0.6 and 4.9 km – consistent with thermal evolution models for the Moon that predict a period of global expansion within the first billion years of lunar history (Solomon and Chaiken 1976; Solomon 1977).

Shortening Structures Wrinkle ridges on the Moon are almost entirely restricted to mare deposits (e.g., Bryan 1973); almost all lunar maria host wrinkle ridges (e.g., Plescia and Golombek 1986). Ridges commonly possess both basin-radial and basin-concentric orientations, although numerous examples are arranged without any preferred orientation. Some early studies suggested that mare ridges reflect the emplacement of magma in the

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subsurface (Strom 1964; Hartmann and Wood 1971), but the overwhelming view is that these landforms are the result of subsidence of the mare lavas themselves (e.g., Bryan 1973; Muehlberger 1974; Ronca 1965; Solomon and Head 1980; Plescia and Golombek 1986; Watters and Johnson 2010), with compressive stresses arising at the basin center balanced by extensional stresses (driving graben formation) at the edges (Solomon and Head 1979). Some examples of wrinkle ridges may have been aided by the vertical motion of a superisostatic mantle plug (Byrne et al. 2015). By definition, therefore, the development of lunar wrinkle ridges postdates both the basins in which the maria ponded and the emplacement of the mare lavas themselves. The presence of wrinkle ridges on even the youngest mare deposits (e.g., Hiesinger et al. 2011) indicates that subsidence of these volcanic units continued after the cessation of the last major phase of effusive volcanism in the late Eratosthenian, such that crustal shortening within the maria continued until at least 1.2 billion years ago (Watters and Johnson 2010). Lobate scarps are the dominant tectonic landform on the lunar farside. Examples of these landforms were recognized early in the exploration of the Moon (e.g., Mattingly et al. 1972) but only with the highest-resolution images available. More recent surveys have identified dozens of these structures, distributed across the body (although the preponderance of scarps on the farside remains) (Watters et al. 2010; Banks et al. 2012). Importantly, lunar lobate scarps are relatively well preserved, indicating that they are comparatively young – perhaps forming within the last billion years or so (Binder and Gunga 1985). This inference, premised on the basis that old, small structures would be unable to retain the crisp morphologies of scarps as they appear today, or would be destroyed entirely, is consistent with observations of scarps crosscutting small craters and with crater size–frequency distributions (Watters and Johnson 2010; van der Bogert et al. 2012). These observations also accord with those same lunar thermal evolution models that predict early lunar expansion (e.g., Solomon 1977), as expansion was probably followed by a sustained

Lunar Tectonism, History of

epoch of global contraction that plausibly continued to as recently as 50 million years ago and may even operate today (Watters and Johnson 2010; Watters et al. 2015). Other mechanisms for driving tectonic deformation have operated throughout lunar history but are expected to be at most secondary contributors. For example, mass wasting has been recognized to operate on the Moon (e.g., Pike 1971), and at least some such activity has likely been facilitated by faults within craters formed within the last billion years (e.g., Xiao et al. 2013). It is also possible that the relaxation of an early tidal bulge could have driven tectonic deformation on the Moon, including thrust faulting at low latitudes at the sub- and anti-Earth points (e.g., Melosh 1980), although no corresponding, predicted extensional tectonism is observed today at the poles (Watters et al. 2010). However, Earth may have played a role in lunar tectonism: the superposition of orbital recession stresses on continued global contraction, augmented by diurnal tidal stresses, could have contributed to the formation of lunar lobate scarps (Watters et al. 2015).

Synthesis Collating the observations described above produces a relatively straightforward history of lunar tectonism. After the formation of a global magma ocean, an early phase of global expansion resulted in the emplacement of a system of major but deep dikes within the Moon. This phase of expansion probably ended within 0.5–1 billion years after the Moon formed and was followed by a protracted period of slow global contraction as the interior cooled by radiating heat into space. This contraction may continue to the present and is the principal agent responsible for the Moon’s population of lobate scarps. Within the first few hundred millions of years, a period of intense bombardment of the Moon by asteroids and comets formed the gigantic impact basins on the lunar nearside and farside, fracturing the lunar lithosphere and providing topographic lows into which expansive basaltic lavas later ponded. The subsidence of these lavas drove the formation of

Lunar Tectonism, History of

lunar wrinkle ridges in the interior of the maria, with bending of the surrounding crust leading to the development of graben. Mare subsidence continued long after the basalts were in place, but major graben formation was inhibited not long after the first set of such structures formed as the globally compressive stress state asserted itself. Stresses from orbital interaction with Earth and gravity-driven processes likely exist even today, but are secondary in magnitude to the tectonic phenomena that shaped the Moon throughout much of the first half of the body’s geological history.

Cross-References ▶ Early Geologic History of the Moon ▶ Evolution, Lunar: from Magma Ocean to Crust Formation ▶ Large-scale Faulting on the Moon ▶ Lunar Core Formation ▶ Lunar Geological Timescale ▶ Volcanic Processes on the Moon

References Ahrens TJ, Rubin AM (1993) Impact-induced tensional failure in rock. J Geophys Res 98:1185–1203 Andrews-Hanna JC et al (2013) Ancient igneous intrusions and early expansion of the Moon revealed by GRAIL gravity gradiometry. Science 339:675–678 Banks ME, Watters TR, Robinson MS, Tornabene LL, Tran T, Ojha L, Willliams NR (2012) Morphometric analysis of small-scale lobate scarps on the Moon using data from the Lunar Reconnaissance Orbiter. J Geophys Res 117:E00H11 Binder AB, Gunga H-C (1985) Young thrust-fault scarps in the highlands: Evidence for an initially totally molten Moon. Icarus 63:421–441. Bryan WB (1973) Wrinkle-ridges as deformed surface crust on ponded mare lava. In: Proceedings of lunar science conference, vol 4, pp 93–106 Byrne PK, Klimczak C, McGovern PJ, Mazarico E, James PB, Neumann GA, Zuber MT, Solomon SC (2015) Deep-seated thrust faults bound the Mare Crisium lunar mascon. Earth Planet Sci Lett 427:183–190 Dombard AJ, Gillis JJ (2001) Testing the viability of topographic relaxation as a mechanism for the formation of lunar floor-fractured craters. J Geophys Res 106:27901–27910

845 Golombek MP (1979) Structural analysis of lunar grabens and the shallow crustal structure of the Moon. J Geophys Res 84:4657–4666 Green DH, Ware NG, Hibberson A, Major A (1971) Experimental petrology of Apollo 12 mare basalts, part 1, sample 12009. Earth Planet Sci Lett 13:85–96 Hall JL, Solomon SC, Head JW (1981) Lunar floorfractured craters: evidence for viscous relaxation of crater topography. J Geophys Res 86:9537–9552 Hartmann WK, Wood CA (1971) Moon: origin and evolution of multiring basins. The Moon 3:3–78 Hiesinger H, Head JW (2006) New views of lunar geoscience: an introduction and overview. Rev Mineral Geochem 60:1–81 Hiesinger H, van der Bogert CH, Reiss D, Robinson MS (2011) Crater size–frequency distribution measurements of Mare Crisium. Lunar Planet Sci 42:2179 Johnson BC et al (2016) Formation of the Orientale lunar multiring basin. Science 354:441–444 Jozwiak LM, Head JW, Zuber MT, Smith DE, Neumann GA (2012) Lunar floor-fractured craters: classification, distribution, origin and implications for magmatism and shallow crustal structure. J Geophys Res 117: E11005 Jozwiak LM, Head JW, Wilson L (2015) Lunar floorfractured craters as magmatic intrusions: geometry, modes of emplacement, associated tectonic and volcanic features, and implications for gravity anomalies. Icarus 248:424–447 Lucchitta BK, Watkins JA (1978) Age of graben systems on the Moon. In: Proceedings of lunar and planetary science conference, 9th Geochim Cosmochim Acta, vol 3, pp 3459–3472 Masursky H et al (1978) Apollo over the Moon: a view from orbit. Scientific and Technical Information Office, National Aeronautics and Space Administration, Washington, DC Mattingly TK, El-Baz F, Laidley RA (1972) Observations and impressions from lunar orbit. Apollo 16 Prel Sci Rep, pp 28-1–28-16. http://adsabs.harvard.edu/abs/ 1972NASSP.315..281M. McGovern PJ, Litherland MM (2011) Lithospheric stress and basaltic magma ascent on the Moon, with implications for large volcanic provinces and edifices. Lunar Planet Sci 42. Abstract 2587 Melosh HJ (1980) Tectonic patterns on a tidally distorted planet. Icarus 43:334–337 Muehlberger WR (1974) Structural history of southeastern Mare Serenitatis and adjacent highlands. In: Proceedings of lunar science conference, 5th Geochim Cosmochim Acta, vol 1, pp 101–110 Mueller K, Golombek MP (2004) Compressional structures on Mars. Annu Rev Earth Planet Sci 32:435–464 Nakamura Y, Lammlein D, Latham G, Ewing M, Dorman J, Press F, Toksöz MN (1973) New seismic data on the state of the deep lunar interior. Science 181:49–51 Pike RJ (1971) Some preliminary interpretations of lunar mass-wasting process from Apollo 10 photography. In:

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846 Analysis of Apollo 10 photography and visual observations, NASA SP, vol 232. Scientific and Technical Information Office, National Aeronautics and Space Administration, Washington, DC, pp 14–20 Platz T, Massironi M, Byrne PK, Hiesinger H (2015) Volcanism and tectonism across the solar system: an overview. Geol Soc Lond Spec Publ 401:1–56 Plescia JB, Golombek MP (1986) Origin of planetary wrinkle ridges based on the study of terrestrial analogs. Geol Soc Am Bull 97:1289–1299 Ronca LB (1965) A geological model of Mare Humorum. Icarus 4:390–395 Schultz PH (1976) Floor-fractured lunar craters. The Moon 15:241–273 Schultz PH, Gault DE (1975) Seismic effects from major basin formations on the Moon and Mercury. The Moon 12:159–177 Scott DH, Diaz JM, Watkins JA (1977) Lunar farside tectonics and volcanism. Proc Lunar Sci Conf 8:1119–1130 Shoemaker EM, Robinson MS, Eliason EM (1994) The south pole region of the Moon as seen by Clementine. Science 266:1851–1854 Solomon SC (1977) The relationship between crustal tectonics and internal evolution in the Moon and Mercury. Phys Earth Planet Inter 15:135–145 Solomon SC (1978) The nature of isostasy on the Moon: how big of a Pratt-fall for Airy methods. Proc Lunar Planet Sci 9:3499–3511 Solomon SC, Chaiken J (1976) Thermal expansion and thermal stress in the Moon and terrestrial planets – clues to early thermal history. Proc Lunar Planet Sci 7:3229–3243 Solomon SC, Head JW (1979) Vertical movement in mare basins: relation to mare emplacement, basin tectonics, and lunar thermal history. J Geophys Res 84:1667–1682 Solomon SC, Head JW (1980) Lunar mascon basins: lava filling, tectonics, and evolution of the lithosphere. Rev Geophys Space Phys 18:107–141 Strom RG (1964) Analysis of lunar lineaments, I: tectonic maps of the Moon. University of Arizona Lunar and Planetary Laboratory. Communications 2:205–216 van der Bogert CH, Hiesinger H, Banks ME, Watters TR, Robinson MS (2012) Derivation of absolute model ages for lunar lobate scarps. Lunar Planet Sci Conf 43. Abstract 1847 Watters TR, Johnson CL (2010) Lunar tectonics. In: Watters TR, Schultz RA (eds) Planetary tectonics. Cambridge University Press, New York, pp 121–182 Watters TR et al (2010) Evidence of recent thrust faulting on the Moon revealed by the lunar reconnaissance orbiter camera. Science 329:936–940 Watters TR, Robinson MS, Collins GC, Banks ME, Daud K, Williams NR, Selvans MM (2015) Global thrust faulting on the Moon and the influence of tidal stresses. Geology 43:851–854 Wichman RW, Schultz PH (1995) Floor-fractured craters in Mare Smythii and west of Oceanus Procellarum: implications of crater modification by viscous relaxation and

Lunar Terrane Tectonics igneous intrusion models. J Geophys Res 100:21201–21218 Wieczorek MA, Jolliff BL, Khan A, Pritchard ME, Weiss BP, Williams JG, Hood LL, Righter K, Neal CR, Shearer CK, McCallum IS, Tompkins S, Hawke BR, Peterson C, Gillis JJ, Bussey B (2006) The constitution and structure of the lunar interior. Rev Mineral Geochem 60:221–364 Wilhelms DE (1987) The geologic history of the Moon. U.S. Government Printing Office, Washington, DC Wood JA, Dickey JS, Marvin UB, Powell BN (1970) Lunar anorthosites and a geophysical model of the Moon. In: Levinson AA (ed) Proceedings of Apollo 11 lunar science conference, vol 1, pp 965–988 Xiao Z, Zeng Z, Ding N, Molaro J (2013) Mass wasting features on the Moon – how active is the lunar surface? Earth Planet Sci Lett 376:1–11

Lunar Terrane Tectonics Jianzhong Liu and Jinzhu Ji Center for Lunar and Planetary Sciences, Institute of Geochemistry, Chinese Academy of Sciences, Guiyang, China

Definition A terrane in geology is a terminology to simply describe a series of related rocks formations or an area having a preponderance of a particular rock or rocks groups, which is separated by faults, and it is a shorthand term for a “tectonostratigraphic terrane.” Since there are lots of similar geology and geochemistry characters in the Earth-Moon system, the conception of terrane can be used to study lunar science, generally called Terrane Tectonics, which means a tectonic unit with distinctive characters and unique geologic history. Thus, the study on the division of lunar terrane tectonics can be contributed to deeply understand the geological evolution of the Moon.

Highland-Mare Dichotomy Since Galileo first peered at the Moon through his homemade telescope in the early 1600s, lunar scientists have divided the lunar surface into two major

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Lunar Terrane Tectonics, Fig. 1 Lunar surface image by NASA’s Lunar Reconnaissance Orbiter: dark parts are mare region and brighter parts are highland region

terranes on the basis of surface smoothness and albedo, highland and mare (Fig. 1). The highlands which have been heavily cratered is generally albedo higher than the maria, which are less cratered (hence younger), darker, and lower in elevation (Whitaker 1978). During the early period of Apollo missions, pre-Soviet scientists divided the lunar surface into two first-order tectonic terrains, the highland tectonic terrain and the mare tectonic terrain, and several subtectonic units (Коэлов and Сулиди-Кондратьев 1967, 1968). This highland-mare dichotomy system served well for a long time. The highlands are anorthositic in composition, which make up about 80% of the Moon’s surface (including virtually all of the farside) and are saturated with impact craters. The maria are basaltic in composition and predominately distributed in the nearside. They cover only about 15% of the total lunar area but 31.2% of the nearside of the Moon (Wilhelms et al. 1987).

FHT-PKT-SPAT Ternary Geochemical Terrane Tectonics As increased studies have been conducted on South Pole-Aitken (SPA) basin, the largest and

oldest confirmed lunar impact basin, it has been proved by evidences from different aspects that SPA basin is a region with geochemical features that differs from highland and mare (Gillis et al. 2004; Jolliff et al. 2000; Lucey et al. 1995). In light of global distribution of FeO (Lucey et al. 1995) and Th (Gillis et al. 2004) derived from Clementine multispectral data and Lunar Prospector gamma-ray data, Jolliff et al. (2000) divided the lunar crust into three distinct provinces whose geochemistry and petrologic history make them geologically unique from each other (Table 1 and Fig. 2): (1) the Procellarum KREEP Terrane (PKT), (2) the Feldspathic Highlands Terrane (FHT), and (3) the South Pole-Aitken Terrane (SPAT). The anorthositic part of the Feldspathic Highlands Terrane (FHT-anorthositic) is corresponding many of the lunar highlands and characterized by low FeO (4.2 wt% on average) and very low Th (0.8 ppm). It is composed of anothosite and related feldspar-rich rocks and represents a pure form of the ancient, primary lunar crust. Jolliff et al. (2000) also identified a related buffer terrane, the outer Feldspathic Highlands Terrane (FHT-O), which is similar to FHT, but contains more FeO (5.5 wt%) and Th (1.5 ppm). They believe that

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Lunar Terrane Tectonics, Table 1 FeO and Th concentrations in crust terranes as exposed at the lunar surface (Jolliff et al. 2000)

FHT-An FHT-O Other mare OM, mixed PKT-nonm PKT-mare PKT-mixed SPAT-inner SPAT-outer

FeO, wt% Mean 4.2 5.5 16.2 8.8 9.0 17.3 10.7 10.1 5.7

s.d. 0.5 1.6 2.3 2.3 2.3 1.8 2.6 2.1 1.1

Th, ppm Mean 0.8 1.5 2.2 1.6 5.2 4.9 4.5 1.9 1.0

% of Area (60 S-60 N) 24.8 34.7 2.8 10.2 2.0 7.6 6.9 5.3 5.7

s.d. 0.3 0.8 0.7 0.8 1.4 1.0 2.0 0.4 0.3

Abbreviations: FHT, Feldspathic Highlands Terrane (An, anorthositic; O, outer, mainly basin-ejecta covered); PKT, Procellarum KREEP Terrane (nonm, nonmare); SPA, South Pole Aitken Terrane; “other mare” refers to regions of mare basalt within the FHT-O, for which individual 5 pixels consist entirely of basalt; “OM, mixed” refers to regions surrounding “other mare” where individual pixels contain both mare and nonmare materials

Lunar Terrane Tectonics, Fig. 2 Three major terrranes of the Moon (Jolliff et al. 2000), see images above: (1) The Feldspathic Highlands Terrane (FHT) which includes its somewhat different outer portion (FHT,O); this terrane has low FeO and Th. (2) The Procellarum KREEP Terrane (PKT), characterized by high Th. (3) South Pole Aitken Terrane (SPA Terrane), which has modest FeO and Th. These do not correspond to the traditional divisions into highlands and mare

FHT, outer (BE, CM)

Near Side

Far Side

FeO (wt.%) 0

5

10

15

Near Side Th, ppm 1 2 4 6 8 10 12

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base image: Luccy et al. (1995)

Far Side Th concentrations from Lunar Prospector data, calibrated to landing site soils (Gillis et al., 2000) (From Jolliff et al., 2000.)

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FHT-O is FHT modified ejecta from huge impact basins and some mare basalt deposits. These highland terranes cover 65% of lunar surface. The Procellarum KREEP Terrane (PKT) dominates the nearside of the Moon. “KREEP” is an acronym for lunar rocks that are high in potassium (K), rare earth elements (REE), and phosphorous (P). PKT is a mixture of assorted rocks, including most of the mare basalts on the Moon, and is characterized by high Th (about 5 ppm by average). The South Pole-Aitken Terrane is located on the southern farside of the Moon. Jolliff et al. (2000) divided the terrane into SPAT-inner and

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SPAT-outer parts. SPAT-inner has moderate FeO (average of 10.1 wt%) and Th (1.9 ppm). SPATouter has less FeO (5.7 wt%) and Th (1.0 ppm).

Highland-Mare-SPA Ternary Terrane Tectonics The three geochemical terrane tectonics systems proposed by Jolliff et al. (2000) supplies with a good model to summarize the geochemical features of the Moon but does not concern the geophysics and lithosphere structure in term of terrane tectonics. GRAIL and LOLA data provide

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Lunar Terrane Tectonics, Fig. 3 Units of lunar tectonic system based on different characters: (a) bouguer gravity anomaly, (b) crust thickness, (c) Th concentration, (d) FeO concentration, (e) the LOLA elevation data, and (f) a

comparison of three lunar tectonic domains’ boundary based on different study objects. a- Mare Tectonic Domain, b- Highland Tectonic Domain, c- South Pole-Aitken Tectonic Domain

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the possibility to investigate the lunar geophysics and crustal in a refined resolution (Wieczorek et al. 2013). Liu et al. (2015) divided the lunar crust into three tectonics units: (1) the Highland Tectonics, (2) the Mare Tectonics, and (3) the South-pole Aitken Basin Tectonics. Furthermore, Guo et al. (2016) combined the study of the characteristics of lunar geophysics, geochemistry, and topography, which provided more evidences to support this lunar terrane tectonic system (Fig. 3). The three terrane tectonics which extend both on the horizontal surface and in the vertical interior are the Mare Tectonic Domain which mainly covers the Procellarum and its neighbor mare basins, the Highland Tectonic Domain which mainly covers the highland in the farside, and the South Pole-Aitken Tectonic Domain which is mainly occupied by the great South Pole-Aitken basin (Fig. 3).

Lunar Transient Phenomena Lucey PG, Taylor GJ, Malaret E (1995) Abundance and distribution of iron on the moon. Science 268(5214):1150–1153 Whitaker EA (1978) Galileo’s lunar observations and the dating of the composition of “sidereus nuncius”. J Hist Astron 9:155–169 Wieczorek MA, Neumann GA, Nimmo F et al (2013) The crust of the moon as seen by grail. Science 339(6120):671–675 Wilhelms DE, Mccauley JF, Trask NJ (1987) The geologic history of the moon. US Government Printing Office, Washington, DC

Lunar Transient Phenomena Anthony Cook Department of Physics, University of Aberystwyth, Aberystwyth, Ceredigion, UK

Definition Cross-References ▶ Estimate of Lunar TiO2 and FeO with M3 Data ▶ Lunar Landscape, Highlands ▶ Lunar Landscape, Maria ▶ Lunar Rocks ▶ Regolith Thickness

References Gillis JJ, Jolliff BL, Korotev RL (2004) Lunar surface geochemistry: global concentrations of Th, K, and Feo as derived from lunar prospector and clementine data. Geochim Cosmochim Acta 68(18):3791–3805 Guo DJ, Liu JZ, Ji JZ et al (2016) Preliminary study on the global geotectonic framework of the moon. Chin J Geophys 59(10):3543–3554 Jolliff BL, Gillis JJ, Haskin LA et al (2000) Major lunar crustal terranes: surface expressions and crust-mantle origins. J Geophys Res-Planets 105(E2):4197–4216 Коэлов ВВ, Сулиди-Кондратьев ЕД (1967) Селенотектоника, Прцроа, No10, ст.47–57 Коэлов ВВ, Сулиди-Кондратьев ЕД (1968) Основные проблемы селенотектоники, Сб. “Пробл. Геоuм. u Космол.” Иэд. “Наука”, М., ст. 234–241 Liu JZ, Guo DJ, Ji JZ et al (2015) Lunar tectonic framework and its evolution inhomogeniety. J Deep Space Explor 2(1):75–79. (in Chinese with English abstract)

A transient lunar phenomenon (or TLP) is a shortterm change on, or above, the lunar surface and can take the form of a colored glow, a brightness variability, an obscuration of detail, gray components to a shadow, or flashes. As far as Earthbased astronomers can tell, no permanent lunar surface changes result, hence why the phenomena are “transient” in nature. In the USA, TLP is sometimes referred to as LTP or lunar transient phenomenon.

Overview The topic of TLP is controversial for three reasons. Firstly, the Moon is essentially a geologically dead world (Heiken et al. 1991), and so astronomers should not expect to see visibly active kilometer-scale surface activity occurring in this modern era. Secondly, the majority of TLPs were discovered by visual Earth-based telescopes which some critics explain away as terrestrial atmospheric, telescope optical, or even psychological factors (Dobbins and Sheehan 2014). Thirdly, TLPs must be extremely rare (Hynek et al. 1976), which makes them difficult to search

Lunar Transient Phenomena

for. Nevertheless, among the nearly 3,000 claims of TLP sightings (O’Connell and Cook 2013), there have been a few well-authenticated cases. It is also possible that at least some flashes seen on the Moon, by visual observers in the past, could be attributed to impact events of the same type that we have modern-day video confirmation of (Cudnik 2010). Several plausible theories, explaining the mechanisms behind TLP, have been published and will be outlined below.

History The majority of TLPs have been catalogued in two NASA publications (Middlehust et al. 1968; Cameron 1978) and one online catalogue extension (Cameron 2006a). Cameron introduces a weighting system for TLP, namely, 1 for a report by an inexperienced observer and up to 5 for a highly authenticated observation of a TLP. While the above are fairly comprehensive catalogues, they do have some typographical errors and probable observational mistakes, and the Cameron catalogues may have over optimistically high weights associated with some of the TLPs (Dobbins and Sheehan 2014). Nevertheless, these catalogues do form a useful starting point for initial studies of this topic. A combined and revised catalogue, with more rigorously assigned weights, is being constructed from the above and also from the archives of the Association of Lunar and Planetary Observers (ALPO) and the British Astronomical Association (BAA), by the Department of Physics at the University of Aberystwyth (Cook et al. 2010). The earliest TLP noted was from a naked-eye sighting of a light on the Moon from around 557 AD (Newton 1972), though a better documented and more topical pre-telescopic sighting in 1178 comes from the writings of Gervase of Canterbury; this has been attributed to the formation of the geologically young, bright ray crater, Giordano Bruno (Hartung 1976); however, recent age estimates for this crater suggest that it may be too old (Basilevsky and Head 2012). In the telescopic age, there have been a number of famous accounts of TLP. For example,

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in the eighteenth century, Sir William Herschel described in some of his observations (Klado 1961) what he referred to as lunar volcanoes seen in Earthshine. It has been argued that these may have been misidentifications of bright ray craters, in particular Aristarchus; however, this does not explain the reddish color seen on 1783 May 4. Another topical debate from the eighteenth century was a reported change in appearance of the crater Linne, but this turned out to be due to a combination of earlier descriptive inaccuracies and map errors (Moore 1977). Of the more notable TLPs of modern times, these have included Kozyrev’s spectra of gas emissions from the central peak of Alphonsus in 1958 (Kozyrev 1962; Kalinyak and Kamionko 1962), the pseudopeak effect seen in the crater Herodotus from the 1950s onwards (Cook and Dobbins 2012), bright flashes seen on the Moon in the 1940/50’s era (e.g., Thornton 1947; Stuart 1957), various claimed observations of lunar luminescence (Kopal and Rackham 1963; Link 1972), the Lowell Observatory sighting of red spots in the Aristarchus area in 1963 (Greenacre 1965; O’Connell and Cook 2013), the 1983 Torricelli B event (Cook 2000), and the Langrenus polarized light events of 1992 (Dollfus 2000).

Observing Programs Several observing programs have spent time looking for TLP. The two NASA-backed ones from the 1960s were Project Moon-Blink (Trident Engineering Associates 1966), organized by Winifred Sawtell Cameron, operating out of NASA’s Goddard Space Flight Center, and another team led by Allen Hynek, from Northwestern University, using the Corralitos Observatory (Hynek et al. 1976) in New Mexico. Both projects utilized electronic imaging cameras behind rotating filter wheels. If a colored area was present on the lunar surface, say red, then through a red filter it would be bright on a monochrome cathode ray tube viewing screen and through a blue filter it would be dark, and the net result would be an obvious blink effect on the viewing screen. The Project Moon-Blink system

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equipment were supplied to 22 observatories with a minimum telescope aperture of 38 cm and detected several TLPs over its lifetime from 1964 to 1966 (Trident Engineering Associates 1966). The Corralitos team detected no TLP, despite putting in over 6,466 h of observing time, between 1966 and 1973 (Hynek et al. 1976). They did however detect very large area “blue clearing” effects on a few occasions, where a UV excess was observed, but this was discounted as a TLP because of the large surface area of the Moon involved and the fact that they noted it would occur close to the full Moon and also when the Moon was at a high altitude above the horizon. However, a full explanation of the “blue clearing effect” was never given. During the run-up to the Apollo missions, a large number of amateur astronomers participated in two further projects: ARGUS-ASTRONET and LION (Schneider 1970). During this time and subsequently, the lunar sections of the Association of Lunar and Planetary Observers (ALPO), the British Astronomical Association (BAA), and the American Lunar Society have continued to monitor the Moon but at a lower level of interest, though there was again some extra support during the Clementine (Buratti et al. 2000) and Lunar Prospector missions (Darling 1998). Mobberley (2013) has questioned the reliability of amateurbased networks, in particular the number of small telescopes used and the problems associated in the influx of inexperienced/overenthusiastic observers. Amateur astronomers still work on TLP projects in 2014 but are more pragmatic, concentrating on disproving past TLPs by reobserving the same sites under similar illumination and where possible similar (topocentric) libration. Their aim is to establish the normal appearance of a lunar formation, and if what was reported for a past TLP repeats, then it was probably not a TLP originally but something like a natural color (McCord 1968), a low-texture area appearing fuzzy (Cook 2013), a sunlit terrain protruding from a shadow (Lena and Cook 2004), etc. The AEOLUS (atmosphere from Earth, orbit, and lunar surface) project, led by Arlin Crotts, built and operates a dual monitoring telescope system operating at Cerro Tololo, Chile, and

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Rutherford Observatory, New York, USA. The system captures the whole nearside lunar disk in white light and compares images taken every 20 s. When a change is found in one telescope, it can be checked for on the second telescope, to make sure that it is not a result of some local effect. It was reported (Crotts et al. 2009) that one month’s worth of continuous observation had been made with a sensitivity to changes at the 1–2 % level. A later publication (Crotts 2010) mentions that 200,000 images had been taken and some plausible optical transients found but does not elaborate on what these might be.

TLP Statistics Middlehurst (1966) was able to show that there was no correlation between TLP and the solar cycle. Chapman (1967) had suggested that there may be a correlation between TLP in Aristarchus and the Earth’s tidal pull, though Cook (2011), using a larger dataset, shows this not to be the case. Middlehurst and Moore (1967) plot the locations of TLP sites and deduce that these tend to be distributed around mare edges, something which is confirmed later by Crotts (2008). Cameron (2006b) investigated many physical parameters which might have been associated with TLP, e.g., magnetopause crossing, perigee, apogee, etc., and deduced that the only one that showed any correlation was that TLPs seem to occur more frequently near the terminator. However, this view is contradicted by Cook et al. (2010) who show after normalizing for observational bias, i.e., where astronomers prefer to look on the Moon, that TLPs occur more frequently toward local noon on the lunar surface. In another statistical analysis, using TLP reports from the Middlehurst et al. (1968) catalogue, Crotts (2008) attempts to remove observational bias that favors observers concentrating on TLP site craters. He did this by comparing pre-1930 and post-1930 TLP reports. 1930 was picked as a division point in time, to avoid overreporting artifacts when TLPs start to become overly interesting to the astronomical community. He was able to show that seven lunar features

Lunar Transient Phenomena

(Aristarchus, Plato, Mare Crisium, Tycho, Kepler, and Copernicus) were statistically significant sites for TLPs and also confirmed the Middlehurst and Moore (1967) finding that TLPs were more likely to be located near mare edges than elsewhere.

Theories Volcanism: Although no longer a tenable theory since the modern era of spaceflight, it was one of the earliest explanations for TLP. This theory became popular after Sir William Herschel reported the presence of lunar volcanoes in Earthshine (Klado 1961). However, with modern hindsight, we know that the last throes of endogenic lunar volcanism were about 1 billion years ago (Ziethe et al. 2009), as determined by crater count age estimates. Interestingly, in the 1960s, Hartman and Harris (1968) suggested that the red glow from a 1963 TLP observation by Greenacre and Barr was due to the incandescent black-body radiation from a fire fountain effect near Aristarchus  though there is no evidence for any resulting annular surface deposit effect seen in modern-era spacecraft images (O’Connell and Cook 2013). Small-scale volcanic flows are still possible on the Moon, via impact melt (Carter et al. 2012); however, with the present low cratering rate, any new craters would be too small to be seen from Earth, and any resulting impact melt incandescent glow is unlikely to be seen either, unless on the nightside of the Moon, and for an impact larger than those observed so far (e.g., Madiedo et al. 2014). Specular Reflection: is another early TLP theory and makes use of the Sun’s glint off of shiny components to rocks on the lunar surface. A variation on the theory involves internal reflection through volcanic glass beads. The net result is that at a specific viewing and illumination angle (equal in the case of reflection), the surface will appear to brighten as the Sun moves through its angular diameter across the lunar sky. Attempts have been made to test this theory on at least three TLP sites: Aristarchus (Cook et al. 2011), Herodotus (Cook and Dobbins 2012), and Torricelli B (Tost 2001), but in all three instances, there

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was no repeat occurrence evidence to support the specular reflection theory. Impacts: are the only instance, so far, of TLP that have been proven (Cudnik 2010). Although all confirmed impact flashes have been seen in Earthshine, one of the brightest (Madiedo et al. 2014) could in theory have been detected against the daylight side of the Moon, and this might explain the Thornton and Stuart flashes (Thornton 1947; Stuart 1957). Unfortunately, impact events do not account for the nonrandom distribution of TLP across the lunar surface (Middlehurst and Moore 1967) nor do the vast majority of impacts account for the typical TLP duration of half an hour (Cameron 2006b). Luminescence: has been proposed to explain some colored TLPs and the observational measurements on the filling in of absorption lines in reflected solar spectra. Early measurements of the latter inferred lunar surface luminous efficiencies of anywhere between 165 nm than nonswirl regions (Hendrix et al. 2016). LAMP measurements showed that swirls have higher FUV ratios at 185.6–147.6 nm compared to the surroundings. Possible explanations include lunar swirls being more immature than young craters, with little to no nanophase iron from weathering and compositional differences.

References Crider DH, Vondrak RR (2003) Space weathering effects on lunar cold trap deposits. J Geophys Res 108. https:// doi.org/10.1029/2002JE002030 Denevi BW, Robinson MS, Boyd AK, Sato H, Hapke BW, Hawke BR (2014) Characterization of space weathering from Lunar Reconnaissance Orbiter Camera ultraviolet observations of the Moon. J Geophys

Lunar Ultraviolet Spectroscopy Res Planets 119:976–997. https://doi.org/10.1002/ 2013JE004527 Feldman PD et al (2014) Upper limits for a lunar dust exosphere from far-ultraviolet spectroscopy by LRO/LAMP. Icarus 233:106–113. https://doi.org/10. 1016/j.icarus.2014.01.039 Gladstone GR, Stern SA, Retherford KD et al (2010) LAMP: the Lyman Alpha Mapping Project on NASA’s Lunar Reconnaissance Orbiter mission. Space Sci Rev 150:161–181. https://doi.org/10.1007/s11214009-9578-6 Gladstone GR et al (2012) Far-ultraviolet reflectance properties of the Moon’s permanently shadowed regions. J Geophys Res 117:E00H04. https://doi.org/10.1029/ 2011JE003913 Grier JA, McEwen AS, Lucey PG, Milazzo M, Strom RG (2001) Optical maturity of ejecta from large rayed lunar craters. J Geophys Res 106(E12):32847–32862. https://doi.org/10.1029/1999JE001160 Hapke B (2001) Space weathering from Mercury to the asteroid belt. J Geophys Res 106:10039–10073. https:// doi.org/10.1029/2000JE001338 Hayne PO et al (2015) Evidence for exposed water ice in the Moon’s south polar regions from Lunar Reconnaissance Orbiter ultraviolet albedo and temperature measurements. Icarus 255:58–69. https://doi.org/10.1016/j. icarus.2015.03.032 Hendrix A, Hansen C, Holsclaw G (2010) The ultraviolet reflectance of Enceladus: implications for surface composition. Icarus 206:608–617. https://doi.org/10.1016/ j.icarus.2009.11.007 Hendrix AR et al (2012a) The lunar far-UV albedo: indicator of hydration and weathering. J Geophys Res 117: E12001. https://doi.org/10.1029/2012JE004252 Hendrix AR et al (2012b) Ultraviolet spectroscopy of the moon: a new look at some not-so-new data. Lunar and planetary science conference 43. https://www.lpi.usra. edu/meetings/lpsc2012/pdf/2839.pdf Hendrix AR et al (2016) Lunar swirls: far-UV characteristics. Icarus 273:68–74. https://doi.org/10.1016/j.icarus. 2016.01.003 Hendrix AR, Hurley DM, Farrell WM, Greenhagen BT, Hayne PO, Retherford KD et al (2019) Diurnally migrating lunar water: evidence from ultraviolet data. Geophys Res Lett 46:2417–2424. https://doi.org/10. 1029/2018GL081821 Henry RC, Fastie WG, Lucke RL et al (1976) A farultraviolet photometer for planetary surface analysis. Moon 15:51–65. https://doi.org/10.1007/BF00562471 Henry RC et al (1995) Ultraviolet albedo of the moon with the Hopkins ultraviolet telescope. Astrophys J 454: L69–L72 Hodges RR (2002) Reanalysis of Lunar Prospector neutron spectrometer observations over the lunar poles. J Geophys Res 107(E12):5125. https://doi.org/10. 1029/2000JE001483 Kramer G et al (2009) The lunar swirls. Lunar and Planetary Institute, Houston

Lunar Volatiles Li S, Milliken RE (2017) Water on the surface of the Moon as seen by the Moon Mineralogy Mapper: distribution, abundance, and origins. Sci Adv 3(9):e1701471. https://doi.org/10.1126/sciadv.1701471 Lucke RL, Henry RC, Fastie WG (1974) Far-ultraviolet lunar mapping from Apollo 17. In: 5th Proceedings of the lunar science conference, pp 469–471 Lucke RL, Henry RC, Fastie WG (1976) Farultraviolet albedo or the moon. Astron J 81:1162–1169 McCoy JE (1976) Photometric studies of light scattering above the lunar terminator from Apollo solar corona photography. Proc Lunar Sci Conf 7:1087–1112. https://ui. adsabs.harvard.edu/abs/1976LPSC....7.1087M/abstract Szalay JR, Horányi M, Colaprete A, Sarantos M (2016) Meteoritic influence on sodium and potassium

863 abundance in the lunar exosphere measured by LADEE. Geophys Res Lett 43:6096–6102. https:// doi.org/10.1002/2016GL069541 Zwinkels J (2016) Light, electromagnetic spectrum. In: Luo MR (ed) Encyclopedia of color science and technology. Springer, New York. https://doi.org/10.1007/ 978-1-4419-8071-7_204

Lunar Volatiles ▶ Lunar Interior, Halogens

L

M

Magmasphere ▶ Lunar Magma Ocean

core from 4.25 to around 1.9 Ga but ceased to exist somewhere 1.92–0.80 Ga. Understanding the origin of the field provides essential information on the nature of the remanent magnetizations.

Magnetic Properties at the Lunar Surface

The Apollo Mission and Natural Remanent Magnetization

Yi-Li Lin McGill University, Montréal, QC, Canada

The Apollo Program starting from 1961 and lasted until 1972 (Cortright 2019) was the first attempt to physically study the moon. The six missions brought back 2196 samples on the lunar surface in total, mainly consists of rocks (66%) and soils (21%) according to Allton (2009). Over 95% of the samples exhibit basaltic liquid somewhere in their ancestries (Gast 1972), suggesting most of the materials can be derived from parent rocks that probably originated as an igneous liquid. The samples brought back from the first landing mission by Apollo 11 revealed magnetic properties (Doell et al. 1970). The work by Nagata and Carleton (1971) showed the magnetizations are mainly due to ferromagnetism of metallic iron, paramagnetism of pyroxenes, and antiferromagnetism of ilmenite. The impact of factors varies among samples. Ferromagnetism is usually due to metallic iron, but sometimes with a few percent of nickel and cobalt (Fuller 1974). Study (Gose et al. 1972) also found an approximately five-time enrichment of native iron in the breccias and soil samples compared to that of the igneous rocks, and that ilmenite is the main metallic conducting

Definition The six Apollo Programs brought back various samples that revealed unexpected magnetic properties, especially the presence of crustal remanent magnetization. Study on the paleointensity shows the field was about 100 mT 3.9 Ga ago and declined around 10 mT 3.2 Ga ago. Currently the Moon does not possess a global magnetic field, but rather magnetized regions distributed nonuniformly on the surface known as magnetic anomalies. Mapping of the magnetic fields with the Magnetometer and Electron Reflectometer equipped on the Lunar Prospector revealed most regions of strong crustal magnetization exist antipodal to the four known impact basins. Theories suggest the origin of the magnetic field are either through thermoremanent magnetization or through shock remanent magnetization. It is only known that the Moon once had a molten dynamo © Springer Nature Switzerland AG 2023 B. Cudnik (ed.), Encyclopedia of Lunar Science, https://doi.org/10.1007/978-3-319-14541-9

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mineral in most of the samples (Strangway et al. 1977). What was even more surprising was the discovery of crustal remanent magnetization (Fuller 1974), especially natural remanent magnetization (NRM) on the lunar sample from Apollo 11 mission (Strangway et al. 1970). The nature and origin of the NRM then became a major interest. It is hard to determine the direction of the NRM relative to the moon’s reference axis because the samples have been contaminated and the positions have been altered (Pearce et al. 1973; Sugiura and Strangway 1980). Even if the samples were left unshocked, the results were still uncertain due to the constituent ferromagnetic minerals are usually in multidomain crystal forms, which have poor magnetic recording properties (Weiss and Tikoo 2014). Samples that were mainly studied were lunar basalts and breccias (Collinson 1993). For lunar basalts, the basaltic rocks contain stable but weak NRM (Fuller 1974). Collinson (1985) shows some samples contain one or more secondary components of NRM, and he questions the nature of the remanent magnetism in lunar rocks. It is unknown if the NRM is driven through thermoremanent magnetization (TRM) substantially unaltered by impacts or through other means; or whether some other magnetization process occurs. Collinson (1983) applies the ThellierThellier technique commonly applied to terrestrial igneous rocks, but results were unsuccessful since the samples change characteristics under heat (Strangway et al. 1977). The few data obtained was supplemented through an anhysteretic remanent magnetization (ARM) method (Stephenson and Collinson 1974) and a normalized method developed by Cisowski et al. (1983). Results showed the magnitude of the field (paleointensity) started from about 100 mT 3.9 billion years (Ga) before decaying to around 10 mT 3.2 Ga ago. Studies were also conducted on the nature of NRM in lunar breccias, which is more magnetized when compared to those of igneous rocks (Strangway et al. 1970). According to studies (Nagata et al. 1971; Gose and Carnes 1973), the lunar soil and breccias samples possessed a

Magnetic Properties at the Lunar Surface

significant amount of superparamagnetic iron. The studies also showed that superparamagnetic particles can acquire strong Viscous Remanent Magnetization (VRM) simply by exposing to a field for a period of time. The excess amount of iron found in soils and breccias were possibly result of surface processes related to impacts (Pearce et al. 1972; Cisowski et al. 1973). Nagata et al. (1971) add that the more strongly impacted microbreccias have stronger and stabler remanent magnetization. It was also found despite a large VRM, some breccias still contain a stable remanent magnetization (Gose et al. 1972) with an intensity of about 10
6 emy/gm (Strangway et al. 1971), about the same as that of an igneous rock. The paleointensity data collected through the Thellier, ARM, and normalized method yield a range of 1–10 mT, greater than the field currently observed (Collinson 1993). Additionally, Coleman et al. (1972) discover definitive area of magnetized areas on the Moon, and the regions are expected to carry magnetization of 10
6 emu/gm, same as the retuned samples, implying the data for the samples can possibly be applied for larger volumes of body in the lunar crust as well. It is also important to note that there are large soft components of NRM surrounding each of the samples (Strangway et al. 1977) which can cause spurious effects when handling without precautions. This superimposed soft component can be removed in fields of 50 Oe or less (Gose et al. 1972). Pearce and Strangway (1972) suggest some of the components of the magnetization were of nonlunar origin, possibly acquired on the return flight. The formation of the component is still unexplained.

Present Study of Magnetic Field Lunar magnetisms have been studied for decades by satellites, ground survey and returned samples (Tsunakawa et al. 2010). The first attempt to measure the Moon’s magnetic properties took place in 1959 by Luna 2 magnetometer (Dolginov et al. 1961), followed by measurements by Luna 10 (Dolginov et al. 1966) and Explorer 35 (Sonett

Magnetic Properties at the Lunar Surface

et al. 1967). Before this, no evidence had suggested that the Moon had a magnetic field in the present nor in the past, according to Collinson (1993). A really weak magnetic field of less than 30 γ was recorded by Luna 2 just 55 km above the lunar surface (Our Geomagnetism Correspondent 1969), and data from Luna 10 and Explorer 35 supported the observation. The detection of local magnetic fields ranging from 6  4 γ to 103  5 γ recorded by magnetometers deployed (Dyal et al. 1972) also provided evidence of weak local fields. Explorer 35 showed that the magnetic moment of the moon is less than 4  1020 cgs units excluding influence of solar winds (Strangway et al. 1970) and limited the overall dipole moment to less than 1020 emu, suggesting a field less than 2 γ (Sonett et al. 1967). Today, it is known that the Moon does not have a global intrinsic magnetic field (less than 2  10
8 of Earth’s dipole moment; Russell et al. 1974), but rather remanent crustal magnetizations distributed widely, weak (Smoluchowski 1973), and nonuniformly across the lunar surface (Coleman et al. 1972; Dyal et al. 1974) known as magnetic anomalies. Study by Anderson (1978) found most magnetized regions are located on the rear side of the moon, expecting somewhere about 75 such regions, and estimated roughly 35–50 regions on the front side, possibly due to lunar magnetic feature less frequent in the maria. No obvious correlation was found between the presence of magnetic fields and lunar surface features (craters, rillies, and mare edges) as well. Currently there are two ways of observing the field: (i) through direct measurements with magnetometers installed on spacecrafts (Hood et al. 2001; Purucker 2008; Tsunakawa et al. 2010), and (ii) through indirect estimate of fields mainly with electron reflectometry (Anderson et al. 1975; Halekas et al. 2001; Mitchell et al. 2008). Harada et al. (2013) propose a new approach with nonadiabatic scattering of plasma sheet electrons mapping small-scale fields (spatial scale less than 20 km). An obstacle while mapping the surface fields is due to a wide spatial scale ranging from a few kilometers to up to hundreds of kilometers. The Lunar Prospector (LP) spacecraft was launched in 1999 (Binder et al. 1998) and the

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Magnetometer/Electron Reflectometer (MAG/ ER) experiment equipped on LP was designed to map crustal magnetic fields over the entire lunar surface (Mitchell et al. 2008) with high sensitivity ~0.01 nT (1 nT ¼ 10
5 G) and high spatial resolution of about 4 km (Lin et al. 1998). The magnetometer (MAG) at the Apollo 11 site reported a field of 36 γ with a gradient less than 4  10
3 γ/cm (Dyal et al. 1970; Dyal and Parkin 1971; Ness 1971). Two field strengths of 43 and 103 γ were reported at the Apollo 14 site 1 km apart (Dyal et al. 1971). Coleman et al. (1972) also discover definite areas of magnetic anomalies with fields typically 1/2 γ by Apollo 15. The evidence combined suggest materials (rocks and soils) with magnetic remanence are distributed over the surface of the Moon (Strangway et al. 1971), since no global field is presented to produce such magnetic susceptibility variations in the samples. Magnetization of lunar crusts was also found by the surface magnetometer (Dyal et al. 1974). Recent studies by LMAG (Kurata et al. 2005; Halekas et al. 2008) suggested a small magnetosphere over the crustal magnetic anomaly, called the mini-magnetosphere for a smaller scale interaction. The MAG measured the vector magnetic fields at the spacecraft by subtracting the external field (Lin et al. 1998) and obtained the lunar crust magnetic field intensity and direction at the spacecraft altitude. The external field was removed through spherical harmonic transform and taking the inverse (Berguig et al. 2013). The product is the internal field, and the vector magnetic crustal anomalies can be determined and studied. On the other hand, the electron reflection magnetometry (ER) revealed hundreds of magnetic patches on the lunar surface (Lin et al. 1998), ranging from 7 km (resolution limit) to about 500 km (Anderson et al. 1975). The ER technique is based on the magnetic mirror effect, the reflection of charged particles from regions of increased magnetic fields. Hood and Artemieva (2008) show the ER mainly provided the mapping of lunar surface intensity. The ER also showed the four known large impact basins (Imbrium, Orientale, Nectaris, Humorum, and Hertzsprung) are demagnetized (Halekas et al. 2001). This is not

M

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surprising, since the current observed magnetism on the lunar surface is confined to the crust, and a large impact would shock and heat the crust above the Curie point to a significant depth. Combining the result by ER with MAG, the direct and indirect measurements were used to map the distribution and intensity of the lunar surface magnetic field at latitudes less than around 30 (Hood 1994). The data by LP found the largest areas of strong crustal magnetization exist antipodal to the impact basins (Lin et al. 1988). This finding can explain some, but not all of lunar crustal magnetism. Hood and Huang (1991) create a model for the origin of magnetic anomalies antipodal to lunar basins, arguing large-scale distribution of lunar crustal magnetization is qualitatively consistent with the impact field hypothesis that will be discussed later. It is generally accepted that the lunar magnetic anomalies originated from natural remanent magnetization of the lunar crust acquired in a certain ambient magnetic field (Tsunakawa et al. 2010). However, the exact origin of lunar magnetic anomalies has been debated. There exist numbers of magnetic anomalies in addition to the strong magnetizations of antipodal regions (Halekas et al. 2001; Hood et al. 2001), and their formations are left unexplained. It is therefore important to understand how the magnetic fields on lunar surface originated.

Investigation on Origin of Lunar Magnetism The origin of lunar magnetic field is still unknown. There is no consensus on the origin since it is difficult to distinguish the nature of most samples (Collinson 1993). Current theories suggest two possible origins: (i) through thermoremanent magnetization (Runcorn et al. 1971) or (ii) through shock remanent magnetization (Hide 1972). Both approaches lack direct evidence that can support either claims even though magnetic measurements have been conducted on lunar surface, according to Harada et al. (2013). What is only known is there exists a dynamo magnetic field from 4.25 Ga (Garrick-Bethell et al. 2017)

Magnetic Properties at the Lunar Surface

to at least 1.9 Ga (Jung et al. 2021), and a magnetic field of about 100 mT between 3.9 and 3.6 Ga (Cisowski et al. 1983). Mighani et al. (2020) conclude the dynamo stops existing around 1.92–0.80 Ga ago by analyzing lunar sample 15498 and 15015 from the Apollo mission. Collinson (1993) suggests that the thermoremanent magnetization (TRM) is mostly consistent with the current observations, with the dynamo generation of the field in the molten, electrically conducting lunar core. Studies (Tikoo et al. 2012; Weiss and Tikoo 2014; Mighani et al. 2020) have confirmed that the moon had a dynamo in periods between around 4.25–1.9 Ga. However, Anderson (1983) shows that there is presently no interior dynamo; and for the moon to have sufficient strength to provide at least some of the magnetism found in the samples, the rotation speed needs to be about 20 times the present speed. In addition, although Stephenson et al. (1975) have shown ancient lunar magnetic field might decrease with time, the paleointensity of the samples had a maximum field in the range of 50–100 mT between 3.9 and 3.6 Ga (Collinson 1993). Objections question the possibility of forming a field this strong during that time, but study by Fuller and Cisowski (1987) show the samples possess ancient magnetic fields with intensities ranging from 0.1 to up to 120 mT between the period 4.0 and 3.5 Ga. The wide range exceeds expectation and may suggest the lunar magnetic field experienced dramatic temporal variations (Weiss and Tikoo 2014). Furthermore, recent studies (Garrick-Bethell et al. 2009, 2017; Weiss and Tikoo 2014) have all strongly suggest the possibility of an ancient lunar dynamo some time in lunar history for at least a billion year long. A long-lasting power source is needed to persist. Tikoo et al. (2017) claims that since no current mechanisms (e.g., core crystallization; Scheinberg et al. 2015 or possession dynamo; Dwyer et al. 2011) are sufficient enough to generate the amount of paleointensity measured in the Apollo samples, it is possible that the dynamo was powered by more than one distinct mechanisms, or the dynamo had switched from a strong-dipole dominating state to a multipolar state after 3.56 Ga (Gastine et al. 2012). However, the exact mechanism is still unknown.

Magnetic Properties at the Lunar Surface

The next plausible origin is through shock remanent magnetization (SRM), or the instantaneous generation of a steady (internal or external) or transient magnetic field during meteorite impacts followed by a shock magnetization of debris (Gattacceca et al. 2010). Crawford and Schultz (1988) claim that the impact-generated magnetic field may account for the thermal magnetic remanence in some of the samples (sample 70019 from Apollo 17) from the Apollo missions, since the age of sample is not in between the predicted span of when the moon had an interior dynamo. Studies have been done through simulation (Hood and Artemieva 2008) and through experiments (Crawford and Schultz 1988; Martelli and Newton 1977) to demonstrate the possibility of impact magnetization as an origin. Studies (Halekas et al. 2003; Mitchell et al. 2008) have relate crustal magnetization with impact structures, and the discovery of NRM in a very young impact glass (after the dynamo ceased to exist) from sample 70019 by Sugiura et al. (1979) all point to the possibility that the field source might be from impact plasmas. However, a most recent study by Jung et al. (2021) states that shock demagnetization alone cannot explain the high paleointensity dispersion between about 3.85 and 3.56 Ga. Another study performed by Oran et al. (2020) with magnetohydrodynamic and impact simulations show that the impact plasma can transiently enhance the field inside the Moon, but the resulting field is too weak to explain the crustal magnetic anomalies observed. Hence the core dynamo theory should account for most of the magnetization on the lunar surface. Until today, the exact origin of the magnetic field on the Moon remains a mystery as well as the cessation of the dynamo, all that is known is there once existed a liquid, electrically conducting interior core which was present around 4.25–1.9 Ga.

Cross-References ▶ Kaguya (SELENE) Mission ▶ Lunar Core Dynamo ▶ Lunar Magnetic Anomalies

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▶ Lunar Mare Basalts, Stratigraphy of ▶ Lunar Surface, Magnetic Field ▶ Surface of the Moon, Distribution of Materials and Structures

References Allton JH (2009) Lunar samples: Apollo collection tools, curation handling, Surveyor III and Soviet Luna samples. NASA, Washington, DC Anderson KA (1978) Distribution of lunar surface magnetic fields. Lunar Planet Sci Conf 9:17–19 Anderson KA (1983) Magnetic dipole moment estimates for an ancient lunar dynamo. J Geophys Res 88:17 Anderson KA, Lin RP, McGuire RE, McCoy JE (1975) Measurement of lunar and planetary magnetic fields by reflection of low energy electrons. Space Sci Instrum 1:439–470 Berguig MC, Hamoudi M, Lemouel JL, Mandea M, Cohen Y (2013) Global mapping of internal lunar magnetic field. Arab J Geosci 6:1063–1072 Binder AB, Feldman WC, Hubbard GS, Konopliv AS et al (1998) Lunar Prospector searches for polar ice, a metallic core, gas release events, and the Moon’s origin. EOS Trans Am Geophys Union 79(8):97–109 Cisowski S, Fuller M, Rose ME, Wasilewski PJ (1973) Magnetic effects of experimental shocking of lunar soil. Lunar Planet Sci Conf Proc 4:3003–3017 Cisowski SM, Collinson DW, Runcorn SK, Stephenson A, Fuller M (1983) A review of lunar paleointensity data and implications for the origin of lunar magnetism. J Geophys Res Solid Earth 88(S02):A691–A704 Coleman PJ, Schubert G, Russell CT et al (1972) Satellite measurements of the moon’s magnetic field: a preliminary report. Moon 4:419–429 Collinson DW (1983) Intensity of NRM. In: Methods in rock magnetism and palaeomagnetism. Springer, Dordrecht Collinson DW (1985) Primary and secondary magnetizations in lunar rocks – implications for the ancient magnetic field of the Moon. Earth Moon Planet 33(1): 31–58 Collinson DW (1993) Magnetism of the Moon – a lunar core dynamo or impact magnetization? Surv Geophys 14:89–118 Cortright EM (2019) Apollo expeditions to the Moon: the NASA history 50th anniversary edition. Courier Dover Publications, Mineola Crawford D, Schultz P (1988) Laboratory observations of impact-generated magnetic fields. Nature 336:50–52 Doell RR, Grommé CS, Thorpe AN, Senftle FE (1970) Magnetic studies of Apollo 11 lunar samples. Geochim Cosmochim Acta Suppl 1:2097–2120 Dolginov SS, Yeroshenko YG, Zhuzgov LN, Pushkov NV (1961) Investigation of the magnetic field of the Moon. Geomagn Aeron 1:18

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870 Dolginov SS, Pushkov NV, Yeroshenko EG, Zhuzgov LN (1966) Measurements of the magnetic field in the vicinity of the moon by the artificial satellite Luna 10, Dokl. Akad. Nauk SSSR, 170, 574–577. Dwyer CA, Stevenson DJ, Nimmo F (2011) A long-lived lunar dynamo driven by continuous mechanical stirring. Nature 479(7372):212–214 Dyal P, Parkin CW (1971) The Apollo 12 magnetometer experiment: Internal lunar properties from transient and steady magnetic field measurements. Proc Second Lunar Sci Conf 3:2391 Dyal P, Parkin CW, Sonett CP (1970) Apollo 12 magnetometer: measurement of a steady magnetic field on the surface of the Moon. Science 169(3947):762–764 Dyal P, Parkin CW, Sonett CP, DuBois RL, Simmons G (1971) Lunar portable magnetometer experiment. In: Apollo 14: preliminary science report, SP-272, 227. NASA, Washington, DC Dyal P, Parkin CW, Cassen P (1972) Surface magnetometer experiments: Internal lunar properties and lunar field interactions with the solar plasma. Lunar Planet Sci Conf Proc 3:2287–2307 Dyal P, Parkin CW, Daily WD (1974) Magnetism and the interior of the Moon. Rev Geophys 12(4):568–591 Fuller M (1974) Lunar magnetism. Rev Geophys 12(1): 23–70 Fuller M, Cisowski SM (1987) Lunar paleomagnetism. Geomatik 2:307–455 Garrick-Bethell I, Weiss BP, Shuster DL, Buz J (2009) Early lunar magnetism. Science 323(5912):356–359 Garrick-Bethell I, Weiss BP, Shuster DL, Tikoo SM, Tremblay MM (2017) Further evidence for early lunar magnetism from troctolite 76535. J Geophys Res Planets 122:76–93 Gast PW (1972) The chemical composition and structure of the Moon. Moon 5:121–148 Gastine T, Duarte T, Wicht J (2012) Dipolar versus multipolar dynamos: the influence of the background density stratification. Astron Astrophys 546:A19 Gattacceca J, Boustie M, Hood L, Cuq-Lelandais J et al (2010) Can the lunar crust be magnetized by shock: experimental groundtruth. Earth Planet Sci Lett 299: 42–53 Gose WA, Carnes JG (1973) The time dependent magnetization of fine-grained iron in lunar breccias. Earth Planet Sci Lett 20(1):100–106 Gose WA, Pearce GW, Strangway DW et al (1972) On the applicability of lunar breccias for paleomagnetic interpretations. Moon 5:106–120 Halekas JS, Mitchell DL, Lin RP, Frey S et al (2001) Mapping of crustal magnetic anomalies on the lunar near side by the Lunar Prospector electron reflectometer. J Geophys Res Planets 106(E11):27841–27852 Halekas JS, Lin RP, Mitchell DL (2003) Magnetic fields of lunar multi-ring impact basins. Meteorit Planet Sci 38: 565–578 Halekas JS, Delory GT, Brain DA, Lin RP, Mitchell DL (2008) Density cavity observed over a strong lunar

Magnetic Properties at the Lunar Surface crustal magnetic anomaly in the solar wind: a minimagnetosphere? Planet Space Sci 56(7):941–946 Harada Y, Machida S, Saito Y, Yokota S et al (2013) Smallscale magnetic fields on the lunar surface inferred from plasma sheet electrons. Geophys Res Lett 40: 3362–3366 Hide R (1972) Comments on the Moon’s magnetism. Moon 4:39 Hood LL (1994) Frozen fields. Earth Moon Planet 67(1–3):131–142 Hood LL, Artemieva NA (2008) Antipodal effects of lunar basin-forming impacts: initial 3D simulations and comparisons with observations. Icarus 193:485–502 Hood LL, Huang Z (1991) Formation of magnetic anomalies antipodal to lunar impact basins: two-dimensional model calculations. J Geophys Res Solid Earth 96(B6):9837–9846 Hood LL, Zakharian A, Halekas J, Mitchell DL et al (2001) Initial mapping and interpretation of lunar crustal magnetic anomalies using Lunar Prospector magnetometer data. J Geophys Res Planets 106(E11): 27825–27839 Jung J, Tikoo S, Gattacceca J, Lepaulard C (2021) Shock demagnetization does not fully explain variations in the lunar paleointensity record. Lunar and Planetary Science Conference, 2020, Houston, United States. hal03532927 Kurata M, Tsunakawa H, Saito Y, Shibuya H, Matsushima M, Shimizu H (2005) Minimagnetosphere over the Reiner Gamma magnetic anomaly region on the Moon. Geophys Res Lett 32(24):24205 Lin RP, Anderson KA, Hood LL (1988) Lunar surface magnetic field concentrations antipodal to young large impact basins. Icarus 74(3):529–541 Lin RP, Mitchell DL, Curtis DW, Anderson KA et al (1998) Lunar surface magnetic fields and their interaction with the solar wind: results from Lunar Prospector. Science 281(5382):1480–1484 Martelli G, Newton G (1977) Hypervelocity cratering and impact magnetisation of basalt. Nature 269:478–480 Mighani S, Wang H, Shuster DL, Borlina CS et al (2020) The end of the lunar dynamo. Sci Adv 6: eaax0883 Mitchell DL, Halekas JS, Lin RP, Frey S et al (2008) Global mapping of lunar crustal magnetic fields by Lunar Prospector. Icarus 194(2):401–409 Nagata T, Carleton BJ (1971) Natural remanent magnetization and viscous magnetization of Apollo 11 lunar materials. J Geomagn Geoelectr 22(4):491–506 Nagata T, Fisher RM, Schwerer FC, Fuller MD, Dunn JR (1971) Magnetic properties and remanent magnetization of Apollo 12 lunar materials and Apollo 11 lunar microbreccia. Lunar Planet Sci Conf Proc 2:2461–2476 Ness NF (1971) Interaction of the solar wind with the Moon. Phys Earth Planet Inter 4(3):197–198 Oran R, Weiss B, Shprits Y, Miljkovic K, Toth G (2020) Was the moon magnetized by impact plasmas? Sci Adv 6(40):eabb1475

Magnetometry Our Geomagnetism Correspondent (1969) Lunar magnetism: is the Moon magnetic? Nature 221:415–416 Pearce GW, Strangway DW (1972) Cause of secondary magnetization in lunar samples. NASA special publication, 315, pp 7–55 Pearce GW, Williams RJ, McKay DS (1972) The magnetic properties and morphology of metallic iron produced by subsolidus reduction of synthetic Apollo 11 composition glasses. Earth Planet Sci Lett 17(1):95–104 Pearce GW, Gose WA, Strangway DW (1973) Magnetic studies on Apollo 15 and 16 lunar samples. Lunar Planet Sci Conf Proc 4:3045 Purucker ME (2008) A global model of the internal magnetic field of the Moon based on Lunar Prospector magnetometer observations. Icarus 197(1):19–23 Runcorn SK, Collinson DW, O’Reilly W, Stephenson A et al (1971) Magnetic properties of Apollo 12 lunar samples. Proc R Soc Lond A Math Phys Sci 325: 157–174 Russell CT, Coleman PJ Jr, Lichtenstein BR, Schubert G (1974) The permanent and induced magnetic dipole moment of the moon. Lunar Planet Sci Conf Proc 5: 2747–2760 Scheinberg A, Soderlund KM, Schubert G (2015) Magnetic field generation in the lunar core: the role of inner core growth. Icarus 254:62–71 Smoluchowski R (1973) Lunar tides and magnetism. Nature 242:516–517 Sonett CP, Colburn DS, Currie RG (1967) The intrinsic magnetic field of the Moon. J Geophys Res 72(21): 5503–5507 Stephenson A, Collinson DW (1974) Lunar magnetic field palaeointensities determined by an anhysteretic remanent magnetization method. Earth Planet Sci Lett 23(2): 220–228 Stephenson A, Runcorn SK, Collinson DW (1975) On changes in the intensity of the ancient lunar magnetic field. Proc Lunar Sci Conf 6:3049–3062 Strangway DW, Larson EE, Pearce GW (1970) Magnetic studies of lunar samples – breccia and fines. Geochim Cosmochim Acta Suppl 1:2435 Strangway DW, Pearce GW, Gose WA, Timme RW (1971) Remanent magnetization of lunar samples. Earth Planet Sci Lett 13(1):43–52 Strangway DW, Pearce GW, Olhoeft GR (1977) Magnetic and dielectric properties of lunar samples (in Russian). In: Kosmochimiya Luny i Planet, edited by A. P. Vinogradov, pp. 712–728 Sugiura N, Strangway DW (1980) Comparisons of magnetic paleointensity methods using a lunar sample. Lunar Planet Sci Conf Proc 11:1801–1813 Sugiura N, Wu YM, Strangway DW, Pearce GW, Taylor LA (1979) A new magnetic paleointensity value for a “young lunar glass”. Lunar Planet Sci Conf Proc 10: 2189–2197 Tikoo SM, Weiss BP, Buz J, Lima EA et al (2012) Magnetic fidelity of lunar samples and implications for an ancient core dynamo. Earth Planet Sci Lett 337:93–103

871 Tikoo SM, Weiss BP, Shuster DL, Suavet C et al (2017) A two-billion-year history for the lunar dynamo. Sci Adv 3:e1700207 Tsunakawa H, Shibuya H, Takahashi F et al (2010) Lunar magnetic field observation and initial global mapping of lunar magnetic anomalies by MAP-LMAG onboard SELENE (Kaguya). Space Sci Rev 154:219–251 Weiss BP, Tikoo SM (2014) The lunar dynamo. Science 346(6214):1246753

Magnetometry D. Waller1 and B. E. Strauss2 1 Johns Hopkins Applied Physics Laboratory, Laurel, MD, USA 2 National Institute of Standards and Technology, Gaithersburg, MD, USA

Definition/Description Magnetometry is the method through which magnetic information is collected using instruments called magnetometers. Magnetometers have existed in simplistic forms for over 2500 years and were primarily used for navigation and warfare until the Cold War and subsequent Space Race, when technological funding shifted to space exploration. Precise sensors were developed to investigate magnetic fields in extreme detail, leading to improved communication and navigation technology and the discovery of magnetic fields throughout the solar system, including on Earth’s natural satellite: the Moon. Lunar magnetometry expanded rapidly during the Space Race but fell off as Cold War tensions dissolved and remained limited to a handful of measurements from orbital and surface magnetometers until a revival of lunar exploration in recent years. Two missions over the past two decades carried magnetometers to the Moon, adding a wealth of data to previous studies and allowing for global studies of magnetism and its relationship to other properties of the Moon. A further improved understanding of the Moon’s magnetic history and current level of magnetism will be delivered by a new generation of

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magnetometers on a promising number of missions planned for the decade ahead.

Introduction For over 2500 years, humans have harnessed the power of Earth’s magnetic field for navigation purposes, using slivers of naturally magnetized stone as compasses for direction finding. In the thirteenth century, Petrus Peregrinus de Maricourt conducted rigorous experiments on magnetism and wrote the first known treatise describing the properties of magnets. Three hundred years later, Sir William Gilbert expanded this work and published his own treatise describing the Earth as a large magnetized sphere. Geophysical theory bloomed into the eighteenth and nineteenth century and mathematicians such as James Maxwell and Carl Gauss developed methods for numerical and analytical representation of electromagnetic fields. Sources of magnetic fields were discovered to include electrical currents in moving fluids, ionized gases, and permanent magnets found naturally occurring in ores like iron, cobalt, nickel, and other rare earth elements or products manufactured from these naturally magnetic elements. In the past century it has been discovered that magnetic fields are present throughout the solar system. The interplanetary magnetic field is an extension of the Sun’s internal magnetic field carried by the supersonic flow of solar wind. Many other bodies in the solar system are associated with magnetic fields of interior, induced, or remanent origin. Many missions have proved that although the Moon lacks an internally generated magnetic field like Earth’s magnetosphere, it does possess regions of weak but significant crustal magnetism. This magnetic information is collected by instruments called magnetometers carried on orbital and surface missions, and our knowledge of lunar magnetism has increased rapidly as magnetometer technology has advanced. Analyses of lunar magnetic data have been used to study properties of the Moon such as:

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electrical conductivity, temperature, and structure of the lunar crust and deep interior; lunar magnetic permeability and iron abundance; lunar surface remanent magnetization origin and interaction with solar wind and external magnetic fields; and the tenuous lunar atmosphere and ionosphere. These studies are improved by the availability of high-fidelity magnetic field information gathered by sensitive and accurate magnetometers. Figure 1 shows magnetic field maps created from magnetometer data collected by the Lunar Prospector spacecraft, which measured not only the total strength of the magnetic field but directional components of the field, allowing for geometric analysis of the field structure. Magnetic fields are vector fields that can be described at any point in space by a magnitude and direction, using any chosen coordinate system. Figure 2 illustrates a scalar and vector quantity represented in a generic three-dimensional coordinate system. Magnetic fields can vary in space and time, as described by Maxwell’s equations. In lunar science to present there are two generic classes of magnetometers, based on the type of information returned: (1) Vector magnetometers produce an output proportional to the strength and direction of the magnetic field, referenced to a principal axis in the sensor. The polarity or sign of the output depends on the direction of the field with respect to the magnetometer sensing axes (i.e., Bx, By, Bz in Fig. 2). Vector measurements are sensitive to spacecraft attitude and contamination from spacecraft-generated magnetic fields. An example of a vector magnetometer is a fluxgate magnetometer, which are the most common sensors used in present space exploration. (2) Scalar magnetometers produce an output proportional to the strength of the total magnetic field (i.e., |B| in Fig. 2) such that the output has magnitude but not direction. Scalar measurements are sensitive to spacecraft-generated magnetic fields but agnostic to spacecraft attitude. An example of a scalar magnetometer is

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Magnetometry, Fig. 1 Global maps of lunar magnetism from scalar measurements (top) and vector component measurements (bottom) (Purucker and Nicholas 2010)

an optically pumped magnetometer, which have been used for calibration of fluxgate magnetometers on several spacecraft. Fluxgate magnetometers and optically pumped magnetometers rely on two different mechanisms to record magnetic fields. Fluxgate magnetometers are based on the change of magnetic reluctance of a ferromagnetic core when it is driven by an alternating current saturating field in the presence of a magnetic field. The driving field is

provided with a primary coil and the changes in the reluctance are measured by a secondary coil (Geyger 1962). Optically pumped magnetometers use the physical phenomenon called the Zeeman effect, which is the shift in energy absorption levels due to interaction of atoms with external magnetic field. The source of atoms is typically a gaseous form of an element such as helium or alkaline metals such as rubidium (Bloom 1962; Smith and Sonett 1976).

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Magnetometry, Fig. 2 A three-dimensional coordinate system with a scalar quantity (magnitude of |B|) described by vector components (directionality of Bx, By, Bz)

Magnetometers are often deployed on relatively long booms to limit magnetic contamination from the spacecraft, although several methods aim to remove contamination by isolating static and dynamic spacecraft fields. Static fields may be characterized before a spacecraft is launched and easily removed from raw data, but dynamic fields require advanced computational modeling of spacecraft fields to reduce contamination (Acuña 2002). In parametrizing spacecraftgenerated fields for modeling, Ness et al. (1971) proposed using two magnetometers on one boom to determine the spacecraft field by measuring the gradient between sensors. This was expanded by Neubauer (1975) to as many as four magnetometers on one boom to determine the spacecraft field. Magnetic fields generated by spacecraft are rarely expressed as a perfectly dipole structure that can be easily characterized, while temporal and spatial offset and variation creates differences between sensors. Separating contaminating fields from raw data collected by multiple sensors is not a simple task; however, additional sensors closer to the spacecraft can be useful in identifying and calibrating spacecraft-generated magnetic fields (Ness et al. 1971; Lepping et al. 1974). Many spacecrafts carry at least two magnetometer sensors to reduce ambiguity in data interpretation and provide redundancy in case of sensor failure.

The majority of vector measurements in space have been made with fluxgate magnetometers, first developed by Aschenbrenner and Goubau (1936) who constructed a sensor core in 1928 using a circular bundle of soft iron florist wire. This was effectively the first ring core sensor, although later refinements and advances were made at the US Naval Ordnance Laboratory (Primdahl 1979). There was a rapid transition to use of new ring core sensors during the 1970s due to their low mass, relatively simple circuitry, and high performance. To present, a majority of lunar magnetic field measurements have been made with ring core magnetometers on spacecrafts, landers, and rovers. Despite the popularity of ring core magnetometers, most spaceflight instruments are currently being built using a dwindling number of ferromagnetic cores which were manufactured up to 1996 using unknown processes by companies and research groups that no longer exist (Strauss et al. 2020). The stockpiles of these legacy components are so depleted that some providers have considered destroying old flight-spare hardware to recover and refurbish the cores for use in new missions. Development of future highperformance next-generation sensors requires the ability to produce new high-quality cores, and several teams are currently tackling the challenge to recover lost manufacturing methods and

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develop new methods and materials to explore new magnetometer designs (Lenz and Edelstein 2006; Díaz-Michelena 2009).

Orbital Lunar Magnetometers In early 1959, the Soviet Luna 1 carried three separate single axis fluxgate magnetometers that passed the Moon at a distance of approximately 6000 km, but magnetic fields could not be determined at such a distance. In late 1959, a second attempt was made with Luna 2 carrying another three-component magnetometer experiment, but during its descent before impact, it did not observe magnetic fields above 55 km altitude (Dolginov et al. 1961). In early 1966, the USSR successfully placed Luna 10 in lunar orbit carrying a triaxial fluxgate magnetometer on a 1.5 m boom and confirmed the existence of weak crustal magnetic fields on the Moon (Dolginov et al. 1966). From late 1966 to early 1974, the USSR launched four more magnetometers on Luna 11, 12, 19, and 22 but due to relatively high orbits and low instrument sensitivity, only upper limits on lunar magnetism could be established (Siddiqi 2002; Balogh 2010). In mid-1967, the US Explorer 35 carried a triaxial magnetometer into lunar orbit and found that the Moon does not have an Earth-like magnetosphere but hinted at the existence of patchy crustal magnetic fields. The magnetometer functioned from launch in July 1967 to spacecraft deactivation in June 1973, during which time the mission provided reference data for the surface magnetometers of the Apollo program (Sonett et al. 1967; Mihalov and Sonett 1968). The US Apollo 15 and 16 missions released Particles and Fields Subsatellites (PFS-1 and PFS2, respectively) which both had a biaxial fluxgate magnetometer, with one sensor parallel to the spin axis and one transverse to the spin axis. Both subsatellites returned data that supported the existence of weak but still significant regions of crustal magnetism on the Moon. PSF-1 operated from 4 August 1971 until 3 February 1972, and PSF-2 operated from 24 April 1972 until 29 May 1972 (Coleman et al. 1972).

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NASA’s Lunar Prospector (LP) carried a triaxial ring core sensor fluxgate magnetometer deployed on a 2.6 m boom. Launched on 7 January 1998, the instrument boom was extended during the journey to the Moon to collect calibration data. The mission was deorbited to impact the surface on 31 July 1999. Lunar Prospector provided data for the first global magnetic field maps of the Moon and uncovered crustal magnetic fields strong enough to form the smallest known magnetospheres in the solar system (Binder 1998; Lin et al. 1998). JAXA’s SELENE/Kaguya (かぐや) carried a triaxial ring core sensor fluxgate magnetometer deployed on a 12 m boom. The instrument boom extended on 31 October 2007 and data collection began on 21 December 2007 until the spacecraft was deorbited to impact the surface on 10 June 2009. SELENE also carried a sensor alignment monitor system called the Sensor Alignment Monitor Coil (SAM-C) that could generate a time-varying artificial magnetic field using biaxial coils on the spacecraft end of the boom. In-orbit calibration was performed twice a month to monitor potential sensor alignment drift and ensure consistent data collection (Takahashi et al. 2009). The Korea Aerospace Research Institute (KARI) of South Korea is developing the Korea Pathfinder Lunar Orbiter (KPLO) as a technology demonstration and resource surveyor for future lunar missions and launch is currently scheduled for August 2022. KPLO will carry a magnetometer instrument composed of three triaxial fluxgate magnetometers on a 1.2 m long boom (Shin et al. 2019).

Surface Lunar Magnetometers The magnetic field experiments conducted by the US Apollo missions began with the Lunar Surface Magnetometer (LSM) installed by Apollo 12 astronauts on 19 November 1969, collecting the first magnetic field measurements from the surface of the Moon. The LSM was part of the Apollo Lunar Surface Experiment Packages (ALSEP) carried on the Apollo 12, 15, and 16 missions. The Lunar Surface Magnetometer for

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Apollo 16 used newly developed ring core sensors from the US Naval Ordnance Laboratory (Dyal and Gordon 1973; Snare 1998). The Lunar Surface Magnetometer was designed to be a continuously operating vector instrument with three fluxgate sensors located at the ends of 100 cm-long orthogonal booms. The sensors were 150 cm away from each other and 75 cm above the surface. An articulating mechanism on each boom could place each sensor parallel to selenocentric X, Y, Z coordinates. This allowed for the calculation of gradients between the sensors in all three directions (Dyal et al. 1970; Dyal and Gordon 1973). On Apollo 14 and 16, the Lunar Portable Magnetometer (LPM) was designed to be a selfcontained and portable instrument easily operated by astronauts. Three orthogonal fluxgate sensors were mounted on top of a 75 cm tall tripod. The magnetometer was connected by a 15 m ribbon cable to a battery and digital displays on the Lunar Roving Vehicle, and the sensors were manually rotated 180 on the tripod by astronauts for calibration (Dyal and Gordon 1973). The Soviet Lunokhod 2 rover carried a fluxgate magnetometer mounted on a 2.5 m long boom at the front of its chassis and obtained magnetic field measurements over its 37 km traverse, but these were more difficult to interpret and the location of the original data is currently unknown (Dolginov et al. 1976; Ness 1979). Astrobotic’s Peregrine Mission 1 was selected in 2019 through the NASA Commercial Lunar Payload Services (CLPS) program. The Peregrine lunar lander will carry a boom-mounted triaxial fluxgate magnetometer provided by NASA Goddard Space Flight Center and is currently scheduled to launch in 2022 (Purucker 2019). The Lunar Magnetotelluric Sounder (LMS) was selected in 2019 as part of the Lunar Interior Temperature and Materials Suite (LITMS) through the NASA Lunar Science and Instrument Technology Payloads (LSITP) program. LMS aims to reveal the structure and composition of the Moon’s mantle by investigating secondary electric and magnetic fields at the lunar surface (Grimm et al. 2020). Southwest Research Institute (SwRI) is using a fluxgate magnetometer as part

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of LMS and LITMS is currently scheduled to be delivered in 2023 by Firefly’s Blue Ghost lunar lander. Lunar Vertex was selected in 2021 through the NASA Payloads and Research Investigations on the Surface of the Moon (PRISM) program as a joint lander and rover suite to explore Reiner Gamma, a distinctive natural phenomenon classified as a lunar swirl. Lunar swirls are enigmatic features on the surface of the Moon that are closely related to strong variations in crustal magnetic fields, also known as magnetic anomalies (Hood and Schubert 1980). The Lunar Vertex rover will carry an onboard magnetometer to investigate magnetic fields at the surface of Reiner Gamma. The Johns Hopkins Applied Physics Laboratory is responsible for the payload and the mission is currently scheduled to launch in the 2024 timeframe.

Toward the Future The science of magnetometry has grown rapidly in the past century and continues toward new heights. The instruments described here are certainly not comprehensive but focus on past and future lunar missions. The next generation of lunar magnetometers launching in this decade will provide a wealth of magnetic information to deepen current understanding of lunar magnetism and improve magnetic field modeling efforts.

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Mantle Bloom AL (1962) Principles of operation of the rubidium vapor magnetometer. Appl Opt 1(1):61–67. https://doi. org/10.1364/AO.1.000061 Coleman P Jr, Schubert G, Russell CT, Sharp LR (1972) The particles and fields subsatellite magnetometer experiment. Apollo 15 preliminary science report, NASA SP-289:22-1–22-9 Díaz-Michelena M (2009) Small magnetic sensors for space applications. Sensors 9(4):2271–2288. https:// doi.org/10.3390/s90402271 Dolginov SS et al (1961) Investigation of the magnetic field of the Moon. Geomagn Aeron 1(1):18–25. https://doi.org/10.2514/3.1589 Dolginov SS et al (1966) Magnetic field measurement in the Moon’s neighborhood at the Luna-10 satellite. Dokl Akad Nauk SSSR 170(3):574–577 Dolginov SS et al (1976) Study of magnetic field, rock magnetization and lunar electrical conductivity in the Bay Le Monnier. Moon 15:3–14 Dyal P, Gordon DI (1973) Lunar surface magnetometers. IEEE Trans Magn 9(3):226–231. https://doi.org/10. 1109/TMAG.1973.1067650 Dyal P, Parkin CW, Sonett CP (1970) Lunar surface magnetometer. IEEE Trans Geosci Electron 8(4):203–215. https://doi.org/10.1109/TGE.1970.271391 Geyger WA (1962) The ring core magnetometer – a new type of second-harmonic flux-gate magnetometer. AIEE Trans 81:65–73. https://doi.org/10.1109/TCE. 1962.6373206 Grimm RE et al (2020) A magnetotelluric sounder to probe terrestrial planet and satellite interiors. Abstract #1568 presented at the 51st lunar and planetary science conference, The Woodlands, 16–20 March 2020 Hood L, Schubert G (1980) Lunar magnetic anomalies and surface optical properties. Science 208:49–51. https:// doi.org/10.1126/science.208.4439.49 Lenz J, Edelstein S (2006) Magnetic sensors and their applications. IEEE Sensors J 6(3):631–649. https:// doi.org/10.1109/JSEN.2006.874493 Lepping RP, Behannon KW, Ness NF (1974) Inflight performance of the dual magnetometer method. Abstract #SS21 presented at the American Geophysical Union 55th annual meeting, Washington, DC, 8–12 April 1974 Lin RP et al (1998) Lunar surface magnetic fields and their interaction with the solar wind: results from Lunar Prospector. Science 281:1480–1484. https://doi.org/ 10.1126/science.281.5382.1480 Mihalov JD, Sonett CP (1968) The cislunar geomagnetic tail gradient in 1967. J Geophys Res 73(21):943–959. https://doi.org/10.1029/JA073i021p06837 Ness NF (1979) The magnetic fields of Mercury, Mars, and Moon. Annu Rev Earth Planet Sci 7:249–288. https:// doi.org/10.1146/annurev.ea.07.050179.001341 Ness NF et al (1971) Use of two magnetometers for magnetic field measurement on a spacecraft. J Geophys Res 76(16 ):3 564– 3573. h ttps://do i.o rg/10 .1 029/ JA076i016p03564

877 Neubauer FM (1975) Optimization of multimagnetometer systems on a spacecraft. J Geophys Res 80(22): 3235–3240. https://doi.org/10.1029/JA080i022p03235 Primdahl F (1979) The fluxgate magnetometer. J Phys E Sci Instrum 12(4):241–253. https://doi.org/10.1088/ 0022-3735/12/4/001 Purucker ME (2019) A vector fluxgate magnetometer to Lacus Mortis on the Moon. Abstract #P33C-04 presented at the American Geophysical Union Fall Meeting 2019, San Francisco, 9–13 December 2019 Purucker ME, Nicholas JB (2010) Global spherical harmonic models of the internal magnetic field of the Moon based on sequential and coestimation approaches. J Geophys Res 115(12):e003650. https:// doi.org/10.1029/2010JE003650 Shin J et al (2019) KMAG: the magnetometer of the Korea Pathfinder Lunar Orbiter (KPLO) mission. Abstract #P31C-3490 presented at the American Geophysical Union Fall Meeting 2019, San Francisco, 9–13 December 2019 Siddiqi AA (2002) Deep space chronicle: a chronology of deep space and planetary probes 1958–2000. Monographs in aerospace history #24, NASA SP-2002-4524 Smith E, Sonett CP (1976) Extraterrestrial magnetic fields: achievements and opportunities. IEEE Trans Geosci Electron 14(3):154–171. https://doi.org/10.1109/TGE. 1976.294447 Snare RC (1998) A history of vector magnetometry in space. In: Measurement techniques in space plasmas: fields, Geophysical monograph 103, pp 101–115 Sonett CP, Colburn DS, Currie RG (1967) The intrinsic magnetic field of the Moon. J Geophys Res 72(21): 5503–5507. https://doi.org/10.1029/JZ072i021p05503 Strauss BE, Purucker ME, Sheppard D (2020) Fluxgate ring cores: the past, present, and future of space-based magnetic sensing. Abstract #GP013-0002 presented at the American Geophysical Union Fall Meeting 2020, held virtually, 1–17 December 2020 Takahashi F et al (2009) In-orbit calibration of the lunar magnetometer onboard SELENE (KAGUYA). Earth Planets Space 61(11):1269–1274. https://doi.org/10. 1186/BF03352979

Mantle Claire McLeod and Aleksandra J. Gawronska Department of Geology and Environmental Earth Science, Miami University, Oxford, OH, USA

Introduction The Earth-Moon system resulted from a giant impact between proto-Earth and a small Mars-

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sized object (Theia). Solidification of the coalesced material resulted in internal differentiation and the establishment of a core-mantle-crust structure (e.g., Smith et al. 1970; Wood et al. 1970; Gagnepain-Beyneix et al. 2006; Weber et al. 2011; Trønnes et al. 2019). As it exists today, the lunar mantle is the likely result of differentiation of a Moon-wide magma ocean, the lunar magma ocean (or LMO; Smith et al. 1970; Wood et al. 1970). The depth of this primordial LMO has been extensively investigated over the past half century with estimates ranging from several 100 km (i.e., a shallow LMO) to scenarios in which the whole Moon was completely molten following the giant impact event (e.g., Minear and Fletcher 1978; Charlier et al. 2018; Steenstra et al. 2020). Despite a lack of consensus regarding depth, most Moon-forming models do at least agree that the LMO existed Moon-wide. These models also generally agree that the onset of LMO solidification was marked by crystallization and settling of dense Mg-rich olivine and orthopyroxene, thus establishing early mafic mantle Mantle, Fig. 1 Summary of LMO processes leading to the establishment of olivine-pyroxene ( ilmenite) -bearing cumulates at depth, an anorthitic, primary flotation crust, and an incompatible trace element-enriched reservoir (urKREEP)

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cumulates at the base of the LMO (Fig. 1, e.g., Charlier et al. 2018; Li et al. 2019; Moriarty III et al. 2021). As cooling and crystallization continued, clinopyroxene and anorthitic plagioclase feldspar became liquidus phases. The relatively denser pyroxene crystals continued sinking, while the relatively less dense feldspar crystals rose (or “floated”) to establish a primordial flotation crust (e.g., Walker and Hays 1977; Warren 1990; Dygert et al. 2017). As solidification progressed, the residual LMO melt became relatively enriched in incompatible elements (e.g., K, REEs, P, Th, and U,) leading to the generation of an urKREEP reservoir (Fig. 1; Warren and Wasson 1979), and KREEP-rich clinopyroxenes. Further differentiation led to the crystallization of trace phosphates, and the formation of oxides, along with dense, late-stage cumulates, including ilmenite-bearing cumulates (IBCs; e.g., Zhao et al. 2019). The establishment of these late-stage cumulates prior to complete LMO solidification is proposed to have then initiated overturn (or “gravitational restructuring”; Moriarty III et al. 2021) of the

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lunar mantle due to density instabilities. This reorganization may have involved as much as 50–70% of the IBCs sinking diapirically toward the earlier-formed mafic mantle cumulates, largely without disturbing the shallow urKREEP reservoir (Zhao et al. 2019). These early LMO processes were later followed by partial melting of the olivine and pyroxene cumulates to produce the younger mare basalts. Collectively, these events have led to the broadly stratified lunar mantle that exists today.

Geophysical Constraints To date, no samples of the deep lunar interior have been unequivocally confirmed to exist either as exposed sections on the Moon, or within the sample collections (Moriarty III et al. 2021; Qian et al. 2021). Knowledge of the Moon’s interior structure is therefore primarily derived from geophysical observations (see summary in Fig. 2). The collection of lunar geophysical data began during the Apollo program through the Apollo Lunar Surface Experiment Package (ALSEP). A variety of datasets were collected both via active source experimentation and passive listening and are summarized in detail by Nunn et al. (2020). Mantle, Fig. 2 Simplified cross section of the lunar interior. Note the presence of shallow and deep moonquakes on the lunar nearside and the largest impact basin in the Solar System, the South PoleAitken Basin, on the lunar farside. (Modified from Wieczorek et al. 2006)

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Instruments deployed included a surface gravimeter, heat-flow probes, retroreflectors, seismometers, and surface magnetometers (e.g., Wieczorek 2009; Garcia et al. 2019). The seismometers yielded the highest resolution data and operated from 1969 to 1977 during which they detected 28 shallow moonquakes, ~7000 deep moonquakes, and ~1800 meteoroid impacts (Wieczorek 2009; Garcia et al. 2019). As summarized in Civilini et al. (2021), there are four, naturally occurring, primary sources of lunar seismicity. These include moonquakes at shallow and deep levels (see Fig. 2), impacts, and thermal events associated with significant changes in temperature between day and night conditions. Nunn et al. (2020) and Garcia et al. (2019) further distinguish impacts into (1) artificial impacts on the lunar surface and (2) meteoroid strikes. Interestingly, at the time of the Apollo missions, scientists did not expect to catalog moonquakes, hence their existence was a discovery in itself (see Nakamura 2015, for a historical summary). From the recent work of Watters et al. (2019), the origin of eight shallow moonquakes detected via ALSEP was attributed to fault activity on young thrust faults, with six of these occurring during times when the Moon was close to its apogee (and thus likely experiencing maximum compressional stress). In

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contrast, deep moonquakes at ~800–1200 km depth were the most commonly detected seismic events during ALSEP. These events have since been correlated to tidal stresses and to an increase in the brittle-ductile transition temperature within the lunar mantle (Kawamura et al. 2017). From data acquired on the lunar surface during the Apollo missions, scientists have also determined that the lunar mantle is probably anhydrous and has temperatures well below the mantle solidus for depths 740 km), seismic velocities increase to ~8.15 km/s and ~4.15 km/s, respectively (Gagnepain-Beyneix et al. 2006). Since the deployment of the Apollo network, scientists have debated the existence of a midmantle discontinuity corresponding to an upper depleted mantle of potential pyroxenite composition, a lower, magnesian-rich, primitive mantle, and a discontinuity at ~1200 km depth (1600  C) where melt may reside (e.g., Nakamura 1983, 2005; Gagnepain-Beyneix et al. 2006; Nimmo et al. 2012; Khan et al. 2014; Wieczorek et al. 2006; Garcia et al. 2019).

Remote Sensing Constraints From high resolution gravity data retrieved via the GRAIL (Gravity Recovery and Interior Laboratory) spacecraft, numerous regions of mass concentration associated with large positive gravity anomalies have been identified. These so-called “mascons” are associated with the Moon’s impact

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basins including those that are, and are not, infilled with basaltic lava (Melosh et al. 2013). From the recent work of Zhao et al. (2021), 3-D inversion of GRAIL data was used to propose that following an impact event and the collapse of a transient crater, lunar mantle material upwelled to fill the crater and establish high-density lithologies beneath the Moon’s basins as observed today. The materials that comprise impact basin rings have the potential to originate from a variety of depths within the lunar interior with Lemelin et al. (2019) recently demonstrating that the innermost rings are often dominated by anorthosite. However, from their detailed mineralogical assessment of data acquired via the SELENE (Kaguya) Multiband Imager, a “mantle component” was also proposed to exist with ultramafic material potentially also present below the single pixel scale (3.6  0.2 Byr (Solomon and Head 1980). Most mascons can be explained by the flexural support of mare basalts within the basins (Solomon and Head 1980), though this does not explain all mascons. Some basins exhibit mascons but lack any mare fill (Neumann et al. 1996), such as Hertzsprung, Korolev, and Mendel-Rydberg. Orientale basin is surrounded by an annulus of thickened, though subisostatic, crust (AndrewsHanna 2013). The flexural uplift of the annulus caused the positive gravity anomalies within the basin center. The uplift exceeded ~2 km, increasing the central gravity anomaly by ~200 mGal, which explains a significant fraction of the Orientale mascon, and is an archetype for nonmare mascons (Andrews-Hanna 2013). A nonmare mascon component has a central gravity anomaly at ~180 mGal relative to the mean anomaly outside the basin, leading to mare thicknesses up to 4 km (Bratt et al. 1985). These nonmare mascons were previously hypothesized to arise from a dynamic rebound of the basin floor and underlying mantle following the impact (Neumann et al. 1996). However, this process would require a sufficiently thick lithosphere to exist preimpact (Andrews-Hanna 2013). Crisium (Fig. 1) does not have an extensive ring structure as seen in Orientale. Most of the mare fill is deep in the central depression (Solomon and Head 1980). The topography of the mare is dominated by an annulus of elevated topography. It sits above uplifted mantle material,

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Mascons

Mascons, Fig. 1 Example of mascon “bull’s-eye” morphology at Mare Crisium. Shown is data of Bouguer gravity degree 60 to 660 from GRAIL GRGM900C gravity model. Scale bar 100 km

which contributes to a mass excess beneath the basin (Byrne et al. 2015). Byrne et al. (2015) have also hypothesized that mare-filled mascon basins share similar tectonic characteristics, mainly by deep, underlying thrust faults. Humorum basin exhibits a ring structure obscured by degradation, with a second basin ring that has major mare deposits concentrations, estimated to be emplaced ~3.6 Byr (Solomon and Head 1980). The Nectaris basin mascon is concentrated in the inner ring depression, within a well-developed ring structure. Imbrium has three major rings: an extensively flooded basin, central peak ring partially exposed, and an asymmetric load in the basin between the second and third rings (Solomon and Head 1980). Smythii is the most ancient of the mascon basins, with mare

lavas concentrated in the central depression. Geologic mapping estimated exposed mare of Imbrian age >3.0 Byr (Solomon and Head 1980). No linear rilles are observed, and mare ridges are concentrated in the northeastern part of the basin. Grimaldi is the smallest of the mascon basins, with an exposed surface of about 3.0 Byr (Boyce 1976). It also exhibits extensive linear rilles, especially to the western edge of Oceanus Procellarum.

Cross-References ▶ GRAIL Mission ▶ Impact Processes on the Moon ▶ Lunar Surface, Gravity Field

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References

MIDAS System Andrews-Hanna JC (2013) The origin of the non-mare mascon gravity anomalies in lunar basins. Icarus 222(1):159–168 Andrews-Hanna JC, Stewart ST (2011) The crustal structure of Orientale and implications for basin formation. In 42nd annual lunar and planetary science conference, No 1608, p 2194 Arkani-Hamed J (1973) Viscosity of the Moon. Moon 6(1): 100–111 Arkani-Hamed J (1998) The lunar mascons revisited. J Geophys Res Planets 103(E2):3709–3739 Boyce JM (1976, April) Ages of flow units in the lunar nearside maria based on Lunar Orbiter IV photographs. In: Lunar and planetary science conference proceedings, vol 7, pp 2717–2728 Bratt SR, Solomon SC, Head JW, Thurber CH (1985) The deep structure of lunar basins: implications for basin formation and modification. J Geophys Res Solid Earth 90(B4):3049–3064 Byrne PK, Klimczak C, McGovern PJ, Mazarico E, James PB, Neumann GA et al (2015) Deep-seated thrust faults bound the Mare Crisium lunar mascon. Earth Planet Sci Lett 427:183–190 Melosh HJ (1978) The tectonics of mascon loading. In: Lunar and planetary science conference proceedings, vol 9, pp 3513–3525 Melosh HJ, Freed AM, Johnson BC, Blair DM, Andrews-Hanna JC, Neumann GA et al (2013) The origin of lunar mascon basins. Science 340(6140): 1552–1555 Muller PM, Sjogren WL (1968) Mascons: lunar mass concentrations. Science 161(3842):680–684 Neumann GA, Zuber MT, Smith DE, Lemoine FG (1996) The lunar crust: global structure and signature of major basins. J Geophys Res Planets 101(E7):16841–16863 Neumann GA, Lemoine FG, Smith DE, Zuber MT (1998) Lunar basins: new evidence from gravity for impact-formed mascons. In: New views of the moon: integrated remotely sensed, geophysical, and sample datasets, p 59 Sjogren WL, Smith JC (1976) Quantitative mass distribution models for Mare Orientale. In: Lunar and planetary science conference proceedings, vol 7, pp 2639–2648 Solomon SC, Head JW (1980) Lunar mascon basins: lava filling, tectonics, and evolution of the lithosphere. Rev Geophys 18(1):107–141 Trowbridge AJ, Johnson BC, Freed AM, Melosh HJ (2020) Why the lunar South Pole-Aitken Basin is not a mascon. Icarus 352:113995 Zhao G, Liu J, Chen B, Kaban MK, Du J (2021) 3-D density structure of the lunar mascon basins revealed by a high-efficient gravity inversion of the GRAIL data. J Geophys Res Planets 126(5):e2021JE006841

José M. Madiedo1 and José L. Ortiz2 1 Facultad de Ciencias Experimentales, Universidad de Huelva, Huelva, Spain 2 Instituto de Astrofísica de Andalucía, CSIC, Granada, Spain

The Moon is continuously being impacted by solid objects of different sizes that move in the interplanetary space at tens of thousands of kilometers per hour. These objects, which are called meteoroids, are produced as a result of the fragmentation and degradation of different celestial bodies, mostly comets and asteroids. The Earth also suffers the impact of these particles of interplanetary matter, which in most cases are completely destroyed in the atmosphere before reaching the ground. But since the Moon has no atmosphere, meteoroids hit the lunar surface at high speed. As a consequence of this, these particles are completely destroyed during these collisions, producing new craters and giving rise to brief flashes of light that can be recorded from Earth by means of telescopes (Fig. 1). In general, these flashes last only a fraction of a second.

MIDAS System, Fig. 1 Example of impact flash produced by the collision of a meteoroid on the lunar surface. The event was recorded in the framework of the MIDAS Survey on December 7, 2013 at 19h31m06s UT

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Besides, impact plumes are produced during the generation of new craters. These plumes are formed by materials that are ejected at high temperature from the impact location. The detection and analysis of impact flashes of meteoroids on the lunar surface is a technique that allows us to obtain very valuable information about the flux of interplanetary matter that impacts the Moon and also provides information on different aspects of the meteoroid complex and on hypervelocity impact physics. Thus, this method uses the Moon as a large detector that informs us, for instance, about the frequency of these collisions. And this, in turn, allows us to estimate the flux of rocks that hit our planet (see, e.g., Ortiz et al. 2006; Madiedo et al. 2014; Suggs et al. 2014). Nevertheless, the results obtained from this technique depend on the so-called luminous efficiency. This parameter, which is often represented by the Greek symbol , is the fraction of the kinetic energy Ek of the impactor that is transformed into luminic energy Er during these impacts: Er ¼ Ek The value of this luminous frequency is not known with enough accuracy, but different studies show that it is of about 2  10
3 (see, e.g., BellotRubio et al. 2000a, b). The first systematic attempts that were made to identify impact flashes due to the collision of large meteoroids on the lunar surface by employing telescopes equipped with highsensitivity CCD cameras date back to 1997 (Ortiz et al. 1999, 2000). This pioneer lunar impact survey was later on called MIDAS (Moon Impacts Detection and Analysis System) (Madiedo et al. 2010, 2014, 2015a, b). Other systematic surveys employing the same observational technique were later on developed by other groups, such as, for instance, the ALPO Survey, which was started by NASA at the MSFC in 2006 (Suggs et al. 2008, 2014). This observational technique is based on the monitoring of the nocturnal (non-illuminated) area of the Moon visible from Earth, since impact flashes occurring on the night side of the Moon can be

MIDAS System, Fig. 2 Typical area on the night side of the Moon monitored by the telescopes operating in the framework of the MIDAS Survey

easily highlighted against the low-illumination background. During the observation, the telescopes are tracked at lunar rate, and the cameras are oriented so that the total nocturnal area monitored on the lunar surface is maximized (Fig. 2). Other phenomena can produce false positives that can be easily confused with lunar impact flashes. For instance, cosmic rays impact very often on the sensor of the camera, giving rise to features that look the same as impact flashes. To discard these and other sources of noise, it is necessary to employ at least two telescopes monitoring the same region of the Moon, since features produced by cosmic rays will be recorded by only one of the devices. Other false positives are also often produced by glints caused by the reflection of sunlight on artificial satellites and space debris. These events can be discarded by employing at least two telescopes located at different observatories. Thus, because of parallax, these glints will appear at different positions on the Moon as seen from each observatory. The main objectives of the MIDAS Survey are the following:

MIDAS System

1. To develop an automated system for the detection of flashes produced by the impact of meteoroids on the lunar surface based on telescopes endowed with high-sensibility cameras. 2. To determine, by means of the analysis of these impact flashes, the value of different parameters employed by impact models. These include the luminous efficiency, the mass distribution of rocks impacting the Moon, the flux of meteoroids that hit the Moon and Earth, and the temperature of the impact plumes generated during these collisions. 3. To coordinate these observations with the results obtained by meteor-observing stations employed to study the interaction of meteoroids with the Earth’s atmosphere. This is fundamental in order to determine which is the most likely meteoroid stream associated to a given impact flash (Madiedo et al. 2015a, b). The first remarkable results derived from this survey were obtained during the activity peak of the Leonid meteor shower in 1999. Thus, during this observational campaign, it was possible to unambiguously detect impact flashes produced by meteoroids associated to this major shower. This, in turn, allowed estimating the value of the luminous efficiency  (Bellot-Rubio et al. 2000a, b). Impact flashes produced by Leonid meteoroids in 1999 were also and serendipitously observed by other groups (see, for instance, Cudnik et al. 2002). Since then, this technique has unequivocally detected impact flashes during the period of maximum activity of several major meteor showers (see, e.g., Ortiz et al. 2000, 2002; Yanagisawa and Kisaichi 2002; Cudnik et al. 2002; Yanagisawa et al. 2006; Suggs et al. 2014; Madiedo et al. 2015a, b), having also identified flashes of sporadic origin (see, e.g., Ortiz et al. 2006; Suggs et al. 2014; Madiedo et al. 2014, 2015b). Other remarkable result obtained in the framework of the MIDAS Survey is the detection of the brightest and longest lunar impact flash recorded to date. This extraordinary event, which lasted 8.3 s, took place on September 11, 2013 with a peak brightness equivalent to that of a magnitude 2.9 star (Madiedo et al. 2014). With a luminosity similar to that of the Pole Star,

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theoretically this flash could have been seen from Earth without any optical instrument. The crater associated to the impact that gave rise to this flash has a diameter of 34 m and was observed by the Lunar Reconnaissance Orbiter (http://lroc. sese.asu.edu/posts/810). The survey was started in 1997 at Instituto de Astrofísica de Andalucía, which is located in Granada (Spain). From there, the survey operated two identical Schmidt-Cassegrain telescopes with a diameter of 0.36 m. Since 2001 the project was stablished at the Sierra Nevada Observatory (Granada, Spain). Besides, some observations were also made by means of a 0.8 m Schmidt telescope located at Calar Alto Observatory plus a 40 cm telescope and 20 cm telescope at Huetor Tájar (Granada, Spain) (Ortiz et al. 2000, 2006). The locations from where the project has been developed have continued changing with time. Thus, in 2008 the two telescopes operating at Sierra Nevada moved to a new observatory in Sevilla, and new telescopes were configured at different observatories. The MIDAS Survey is currently being developed from three astronomical observatories located in Spain: Sevilla, La Hita, and La Sagra. Besides, some specific observational campaigns are also being developed at the Astronomical Observatory of Calar Alto. The geographical coordinates of the sites currently involved in the development of this survey are shown in Table. 1. At Sevilla the MIDAS Survey operates two identical 0.36 m Schmidt-Cassegrain telescopes that monitor the same area of the Moon, but also three smaller Schmidt-Cassegrain telescopes. Two of them have a diameter of 0.28 m, and the third one has a diameter of 0.24 m. These telescopes are equipped with f/3.3 MIDAS System, Table 1 Geographical coordinates of the astronomical observatories implied in the MIDAS Survey Observatory name Sevilla La Hita La Sagra Calar Alto

Longitude 5 580 5000 W 3 110 0000 W 2 330 5200 W 2 320 4600 W

Latitude (N) 37 200 4600 39 340 0600 37 590 0200 37 130 2500

Altitude (m) 28 674 1520 2168

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focal reducers to increase the monitored area on the lunar surface and also with monochrome 8-bit high-sensitivity CCD video cameras (model 902H Ultimate, manufactured by Watec Corporation, Japan). These cameras produce analogue interlaced images with a resolution of 720  576 pixels and a frame rate of 25 frames per second. Since December 2014 some of these telescopes also employ optical filters in order to perform observations in different spectral bands. This makes possible to estimate the temperature of MIDAS System, Fig. 3 The 0.4 m Newtonian telescope operating at La Hita Astronomical Observatory in the framework of the MIDAS Survey. (Image credit: Leonor Ana Hernández)

MIDAS System, Fig. 4 In the front are the two 0.36 m telescopes operating at La Sagra Astronomical Observatory in the framework of the MIDAS Survey

MIDAS System

impact plumes produced during the collisions that give rise to lunar impact flashes (Madiedo and Ortiz 2016). In 2013, a 0.4 m Newtonian telescope was installed at La Hita Astronomical Observatory, in Central Spain. This also employs a Watec 902H Ultimate camera (Fig. 3). Two additional 0.36 m Schmidt-Cassegrain telescopes started regular operation at La Sagra Astronomical Observatory in 2015 (Fig. 4). These are equipped with 12-bit digital video cameras (models ASI174MM, manufactured by ZWO,

Modeling of the Lunar Magma Ocean

and QHY174M, manufactured by QHYCCD). Both of these cameras provide a resolution of 1920  1200 pixels. For every telescope, recorder flashes are stored in a database. These events are then examined to determine which of them are false positives and which are produced by impactors hitting the Moon. The resulting lunar impact flashes are then analyzed by means of a dedicated software (Madiedo et al. 2015b).

Cross-References ▶ Lunar Impact Flashes, Causes and Detection ▶ Lunar Impact Event: The 11 September 2013 ▶ Lunar Impact Events, Luminous Efficiency and Energy of

References Bellot-Rubio LR, Ortiz JL, Sada PV (2000a) Luminous efficiency in hypervelocity impacts from the 1999 lunar Leonids. Astrophys J 542:L65–L68 Bellot-Rubio LR, Ortiz JL, Sada PV (2000b) Observation and interpretation of meteoroid impact flashes on the Moon. Earth Moon Planet 82/83:575–598 Cudnik BM, Dunham DW, Palmer DM, Cook AC, Venable RJ, Gural PS (2002) Ground-based observations of high velocity impacts on the Moon’s surface – the lunar Leonid phenomena of 1999 and 2001. Page abstract no.1329 of: 33rd annual lunar and planetary science conference Madiedo JM, Ortiz JL (2016) Measuring the temperature of impact plumes from the analysis of lunar impact flashes. https://www.cosmos.esa.int/documents/ 653713/1000954/08_ORAL_Madiedo.pdf/d64b232 5-8a37-432a-98f1-28784da94a40 Madiedo JM, Trigo-Rodriguez JM, Ortiz JL, Morales N (2010) Robotic systems for meteor observing and Moon impact flashes detection in Spain. Adv Astron 2010:1–5, 167494 Madiedo JM, Ortiz JL, Morales N, Cabrera-Cano J (2014) A large lunar impact blast on 2013 September 11. Mon Not R Astron Soc 439:2364–2369 Madiedo JM, Ortiz JL, Organero F et al (2015a) Analysis of moon impact flashes detected during the 2012 and 2013 Perseids. Astron Astrophys 577:A118 Madiedo JM, Ortiz JL, Morales N, Cabrera-Cano J (2015b) MIDAS: software for the detection and analysis of lunar impact ashes. Planet Space Sci 111:105–115 Ortiz JL, Aceituno FJ, Aceituno J (1999) A search for meteoritic flashes on the Moon. Astron Astrophys 343:L57–L60

901 Ortiz JL, Aceituno FJ, Aceituno J (2000) Optical detection of meteoroidal impacts on the Moon. Nature 405:921–923 Ortiz JL, Quesada JA, Aceituno J et al (2002) Observation and interpretation of Leonid impact flashes on the Moon in 2001. Astrophys J 576:567–573 Ortiz JL, Aceituno FJ, Quesada JA, Aceituno J, Fernández J, Santos-Sanz P, Trigo-Rodriguez JM, Llorca J, Martin-Torres FJ, Montañes-Rodriguez P, Pallé E (2006) Detection of sporadic impact ashes on the Moon: implications for the luminous efficiency of hypervelocity impacts and derived terrestrial impact rates. Icarus 184:319–326 Suggs RM, Cooke WJ, Suggs RJ, Swift WR, Hollon N (2008) The NASA lunar impact monitoring program. Earth Moon Planet 102:293–298 Suggs RM, Moser DE, Cooke WJ, Suggs RJ (2014) The flux of kilogram-sized meteoroids from lunar impact monitoring. Icarus 238:23–26 Yanagisawa M, Kisaichi N (2002) Lightcurves of 1999 Leonid impact flashes on the Moon. Icarus 159:31–38 Yanagisawa M, Ohnishi K, Takamura Y, Masuda H, Sakai Y, Ida M, Adachi M, Ishida M (2006) The first confirmed Perseid lunar impact ash. Icarus 182:489–495

Modeling of the Lunar Magma Ocean Saira Hamid and Joseph G. O’Rourke School of Earth and Space Exploration, Arizona State University, Tempe, AZ, USA

Definition/Description Modeling of the lunar magma ocean (LMO) refers to studies of how the Moon evolved from a putative early state when it was at least partially molten. Simple models were first proposed to explain the anorthositic composition of the highland crust evidenced by Apollo samples. Thermodynamic phase relationships that describe the sequence of minerals produced during fractional and/or equilibrium crystallization are foundational to these models. Matching the observed thickness and purity of the crust requires certain choices of key model parameters: the rheology and bulk composition of the magma ocean, its initial thickness, and the amount of interstitial fluid trapped in

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solidified material. Tidal heating in models can extend the timescales of solidification to match the radiometric ages of crustal samples. Onedimensional “onion shell” models obviously cannot reproduce the crustal dichotomy between the near and far sides of the Moon. However, hemispheric differences during the solidification of the magma ocean can arise in higher-dimensional models from asymmetric impacts, tidal heating, and/or Earthshine. Ultimately, models of the lunar magma ocean become more complex as more is learned about the Moon. Future investigations of the structure, composition, and chronology of the lunar surface and interior will motivate even more progress.

Introduction The Moon was likely at least partially molten after it accreted. A giant impact with proto-Earth is the leading model for the formation of the Moon (e.g., Hartmann and Davis 1975). The classical scenario – a Mars-sized impactor colliding with proto-Earth – involves enough kinetic energy to melt most of the silicates that form the Moon (e.g., Nakajima and Stevenson 2014). Recently, studies have proposed impacts involving even higher amounts of kinetic energy, including two halfEarth-mass impactors (Canup 2012) and small impacts onto a rapidly rotating proto-Earth (Ćuk and Stewart 2012; Lock et al. 2018). Models of the Moon-forming impact are debated because matching all the geochemical constraints is difficult (e.g., Barr 2016). Formation of a LMO is perhaps the least controversial outcome of the Moon’s formation. The range of lithologies seen on the lunar surface also motivated the first magma ocean models. Crystallization of the LMO can produce at least three important types of material (e.g., Wieczorek et al. 2006; Elardo et al. 2011; Charlier et al. 2018). First, mafic cumulates (e.g., olivine and pyroxene) form. These solids are denser than the magma ocean and thus tend to sink toward the bottom. Once most (~68–81 vol.%) of the LMO

Modeling of the Lunar Magma Ocean

has solidified, anorthositic plagioclase begins to crystallize. Anorthosite is less dense than the residual liquid and thus tends to rise to form the lunar crust. Finally, the last residium (>90–95 vol. % solidified) of the LMO is “ur-KREEP,” enriched in potassium (K), rare earth elements (REE), and phosphorus (P) (Warren and Wasson 1979; Shearer et al. 2006) – and also in other incompatible elements. Ilmenite-rich cumulates can also form alongside ur-KREEP. The lunar mantle was likely gravitationally unstable after the LMO fully solidified because ur-KREEP and ilmenite are denser than other crystallization products. Ultimately, models of the LMO are important because they establish the compositional diversity of the Moon’s interior and surface. Models of the LMO share many common features (Fig. 1). Initially, the LMO is hot – above its liquidus temperature. The LMO cools over time as heat is lost from the surface to space – fast enough for vigorous convection to maintain an adiabatic temperature profile. Once the liquidus is crossed, the different types of minerals described above start to crystallize and, eventually, sink or rise. The model has served its purpose once the LMO cools to its solidus temperature, except perhaps locally in some source regions for post-LMO magmatism. Other models, beyond the scope of this topic entry, are required to describe the many interesting dynamics that occur thereafter. For example, mantle overturn after LMO solidification could concentrate KREEP material above the core-mantle boundary and produce a basal magma ocean at later times (Scheinberg et al. 2018). In general, choices that are made when setting up models of the LMO can strongly influence predictions for how the Moon evolves over billions of years. This topic entry lists many of the key parameters that influence the output of models of the LMO. It briefly describes how these outputs relate to some of the outstanding science questions about the formation and evolution of the Moon. Future missions and investigations that would improve models of the LMO are advertised at the end.

Modeling of the Lunar Magma Ocean

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Modeling of the Lunar Magma Ocean, Fig. 1 Models of the LMO must include many processes. Fractional crystallization produces anorthositic crust (a) and mafic cumulates (b), possibly with trapped interstitial fluid. Key model parameters determine if convection in the LMO is vigorous enough to entrain crystals, leading to more equilibrium crystallization, (right). Alternatively, solid minerals may efficiently settle out and become stratified based on their relative densities (left). Once the flotation crust forms,

thermal conduction through that lid controls the cooling rate of the LMO. Impacts can increase the cooling rate (by puncturing the lid) and/or add heat to the LMO (via kinetic energy). Models can also include various sources of internal heating, including tidal dissipation, that extend the lifetime of the LMO. The temperature at the base of the LMO is near the solidus of the mafic cumulates at the relevant pressure. (Adapted from Fig. 3 in Dygert et al. (2017))

Key Model Parameters

(Fig. 1). Smaller crystals could remain entrained by convection. A popular formula from Solomatov et al. (1993) calculates the critical diameter for crystal entrainment:

Crystallization Sequence Fractional crystallization occurs in magma if minerals are removed from the melt once they solidify. In contrast, equilibrium crystallization takes place if solids remain mixed and in equilibrium with the remaining liquid. The existence of the highland crust – made of buoyant plagioclase – means that fractional crystallization dominated the final stages of LMO solidification. Models typically feature one of two end member scenarios: (1) fractional crystallization from the start (e.g., Elkins-Tanton et al. 2011) or (2) a “twostage model” with a period of equilibrium crystallization followed by fractional crystallization once ~50% of the LMO solidified (e.g., Tonks and Melosh 1990; Snyder et al. 1992; Elardo et al. 2011). Fluid dynamics determines the form of crystallization in models of the LMO. If solid crystals are so large that they separate vertically from the magma, then fractional crystallization occurs

1 D¼ Drg

rffiffiffiffiffiffiffiffiffiffiffi aQ , 10CP

ð1Þ

where Δr (kg/m3) is the density difference between solid crystals and magma, g (m/s2) is gravitational acceleration, (Pa s) is the dynamic viscosity of the magma, α(1/K) is the thermal expansivity of the magma, Q (W/m2) is the heat flux from the convecting magma to the surface, and Cp (J/kg/K) is the heat capacity of the magma. Experiments to quantify the rheology of the LMO are valuable because viscosity is very sensitive to temperature and composition (e.g., Dygert et al. 2017). Overall, the takeaway message is that bigger crystals can be entrained in magma that is relatively viscous and/or vigorously convecting (i.e., with higher heat flow).

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Interstitial Melt When minerals crystallize from the LMO, some liquid can be trapped in pore spaces (Fig. 1). The volume fraction of interstitial fluid is an important model parameter from a geophysical and a geochemical perspective. First, the density contrast between solids and the liquid magma governs whether crystals settle (Eq. 1) and the gravitational stability of the mantle after LMO has cooled. Second, trapped liquid can change the bulk composition of the anorthositic crust, the urKREEP material, and the mafic crust. For example, the extremely high plagioclase content of the crust is difficult to explain in models unless processes such as Darcy flow with compaction can squeeze out interstitial melt (e.g., Piskorz and Stevenson 2014). In models, the volume fraction of interstitial melt has been treated as a free parameter with values from ~0% to 10% (e.g., Elkins-Tanton et al. 2011). Recent studies favor smaller amounts of trapped liquids (e.g., Charlier et al. 2018). The observed partitioning behavior of trace elements may imply that the source regions of the mare basalts crystallized with 5 wt. %), and Th (approximately >8 ppm; Giguere et al. 2000; Lawrence et al. 2002; Jolliff et al. 2000; Gillis et al. 2003, 2004), as summarized in Fig. 2. The elevated abundances of these elements indicate that radioactive isotopes may be concentrated within the PKT and could have acted as a heat source within the lunar mantle, leading to partial melting and the production of the basaltic maria (e.g., Haskin et al. 2000; Hess and Parmentier 2001; Shearer et al. 2006). Based on the localized nature of the lunar basaltic maria within the PKT, the KREEPy region within the lunar mantle may only exist in the region directly underlying the PKT (e.g., Haskin et al. 2000). While some authors have argued that Oceanus Procellarum may be an ancient impact structure, the identification of distinct ring structures and corresponding geochemical signatures of such a structure is not clear (see summary in Hiesinger et al. 2003). Visually, this region is

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Moon, Fig. 2 Global maps of the lunar surface. Panels B-E have the same latitude and longitude extent as panel A. (A) Lunar Reconnaissance Orbiter Camera (LROC) Wide Angle Camera (WAC) mosaic (100 m/pixel) of the lunar surface centered on the lunar nearside, available on Quickmap (https://quickmap.lroc.asu.edu/). (B) Global digital elevation model (DEM) from the Lunar Orbiter Laser Altimeter (LOLA) available through NASA Moon Trek (https://trek.nasa.gov/moon). (C) Map of the distribution of Th (in ppm) across the lunar surface as measured by the Lunar Prospector Gamma Ray Spectrometer. Map

available through NASA Moon Trek (https://trek.nasa. gov/moon); for more information, refer to Lawrence et al. 2002. (D) Map of the distribution of FeO (in wt. %) across the lunar surface as measured by Kaguya/SELENE. Map available through the Quickmap tool (https://quickmap. lroc.asu.edu/); for more information, refer to Lemelin et al. 2019. (E) Map of the distribution of TiO2 (in wt. %) across the lunar surface estimated from UV/Vis reflectance. Map available through the Quickmap tool (https://quickmap. lroc.asu.edu/); for more information, refer to Sato et al. 2017

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Moon, Fig. 3 The three major surface terranes as outlined by Jolliff et al. (2000) based on the observations summarized in Fig. 2. Map is centered on the lunar nearside; PKT indicates the Procellarum KREEP Terrane; FHT indicates the Ferroan Highlands Terrane; SPAT indicates the South

Pole-Aitken Terrane. Lines are overlain on a Lunar Reconnaissance Orbiter (LRO) Wide Angle Camera (WAC) mosaic (100 m/pixel) available on Quickmap (https:// quickmap.lroc.asu.edu/)

made up of both light (i.e., anorthositic) and dark (i.e., mare) materials (Jolliff et al. 2000). For a detailed summary of the materials existing within the PKT and their relationship to observed geomorphologic units, see Haskin et al. (2000).

between 3.9 and 1.2 Ga, with the youngest basalts on the lunar nearside appearing to be located within Oceanus Procellarum in the PKT (Hiesinger et al. 2003, 2011; Snape et al. 2019). The greatest volume of basaltic material is estimated to have erupted between 3.7 and 3.3 Ga (Fig. 1; Hiesinger et al. 2003; Hiesinger et al. 2011), at which point degassing via volcanism may have been significant enough to sustain a transient lunar atmosphere (e.g., Needham and Kring 2017; Head and Wilson 2020). Particles within this atmosphere could have settled around the Moon, and may now be preserved within permanently shadowed craters (PSRs) at the lunar poles (e.g., Anand et al. 2012). Volcanic gases may therefore be a component of volatile ices that will likely be sampled at the lunar south pole during future missions to this region (e.g., by the Volatiles Investigating Polar Exploration Rover or VIPER, and the Artemis missions). For further discussion of the different volcanic lithologies that exist on the Moon, including information regarding the primitive lunar glass beads, see Shearer et al. (2006) and Wieczorek et al. (2006).

The Lunar Basaltic Maria

The lunar basaltic maria are not considered a specific terrane in the model of Jolliff et al. (2000), but rather are incorporated into the region in which they exist. The lunar maria are primarily located within the PKT and, considering the thermal and magmatic history of the Moon postLMO, they define a geologically important lunar lithology (see Fig. 4). The basaltic maria overall have been observed remotely to range in bulk wt. % TiO2 from 60 km, see Fig. 3; Wieczorek et al. 2013) and has a relatively high elevation (see Fig. 2; Smith et al. 2010). The other portion corresponds to the remaining crust that is dominated by mafic ejecta deposited by impact processes as a result of basin formation and maria degradation. See Jolliff et al. (2000) for detailed descriptions of both. In general, the lunar anorthositic crust is composed of plagioclase feldspar cumulates (approximately >95% plagioclase feldspar, Shearer et al. 2006) with a minor mafic mineral component. Crustal lithologies found on the Moon also include troctolite (feldspar with olivine), norite (feldspar with orthopyroxene), and pure anorthosite (>98% plagioclase feldspar) (e.g., Hawke et al. 2003; Ohtake et al. 2009; Taylor 2009; Lemelin et al. 2019). Stratigraphically, the upper layers of the lunar crust are characterized by feldspathic lithologies, while the

lower layers are more noritic in nature (Hawke et al. 2003, and references therein). At present, debate exists regarding the crystallization ages of ferroan anorthosite samples which represent the crust, and therefore the overall timing of lunar crust formation. Age estimates range from as old as ~4.5 Ga to as young as ~4.3 Ga (Stöffler et al. 2006; Snape et al. 2016). For further discussion of this, the reader is referred to Elkins-Tanton et al. (2011) and Tartèse et al. (2019). The Plutonic Suites

A number of plutonic suites have been discovered in the lunar sample collection that are interpreted to represent differentiated magmas which intruded into the anorthositic crust. Lunar plutonic samples are generally associated with either the magnesian suite or the alkali suite (e.g., Shearer et al. 2006; Wieczorek et al. 2006; Taylor 2009). For a full summary of the evolved rocks that have been observed and sampled at the lunar surface,

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and their compositional characteristics, see Wieczorek et al. (2006) and Shearer et al. (2006). Specifically, lithologies associated with the magnesian suite contain mafic minerals which have Mg/(Mg þ Fe) values of 0.6 to 0.95 and are the oldest of the plutonic samples at >4.2 Ga in age (Wieczorek et al. 2006). The alkali suite comprises a wide variety of lithologies (e.g., anorthosite, norite, gabbro, felsite, and monzogabbro) that are enriched in alkalis relative to other lunar rocks, but that otherwise vary in major and trace element content. In general, the alkali suite rocks have lower Mg/(Mg þ Fe) values than the magnesian suite, ranging from approximately 0.4 to 0.6, and have noticeably high Na2O (wt. %) and Eu (ppm) contents relative to other ferroan anorthosite samples. For a detailed summary of alkali suite rocks, see Wieczorek et al. (2006). The alkali suite samples are also slightly younger in age, ranging from 4.4 to 4.0 Ga (Shearer et al. 2006). It has been hypothesized that KREEP played an important role during the petrogenesis of both of these plutonic sample suites, ultimately leading to the geochemical enrichments observed (e.g., Jolliff et al. 2000; Shearer et al. 2006; Wieczorek et al. 2006; Taylor 2009). South Pole-Aitken Terrane The third terrane outlined by Jolliff et al. (2000), as determined by remote sensing data (see Fig. 2; Lucey et al. 1998, 2000; Lawrence et al. 2002; Gillis et al. 2003, 2004; Sato et al. 2017; Lemelin et al. 2019), is the South Pole-Aitken Terrane (SPAT; Fig. 3). This terrane outlines an area on the southern lunar farside, where average FeO contents are elevated (up to ~15 wt. %) (e.g., Lucey et al. 1998; Jolliff et al. 2000; Lemelin et al. 2019). The SPAT is on average 8 km deeper than the surrounding feldspathic terrane (e.g., Fig. 2; Smith et al. 2010; Wieczorek et al. 2013). While the SPAT is likely an ancient basin formed via impact, the lunar crust here is still thicker than average; hence, it is unlikely that the associated basin-forming impact exposed mantle materials (Borst et al. 2012; Wieczorek et al. 2013). In confirmation of this, Moriarty III and Pieters (2018) noted that significant exposures of olivine,

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which would otherwise be consistent with derivation from the lunar mantle beneath SPAT, were absent. Jolliff et al. (2000) suggested that this may be either because this terrane was originally part of the thick FHT crust, or because the area is covered with an impact melt sheet. The mineralogy and stratigraphy of the SPAT are both consistent with a melt sheet interpretation (e.g., Pieters et al. 2001; Hurwitz and Kring 2014; Vaughan and Head 2014). Specifically, the SPAT is mineralogically dominated by pyroxene. At its center, these pyroxenes are high in Ca and Fe, and broadly represent basaltic-type rocks, while the remainder of SPAT contains Mg-rich pyroxenes consistent with more noritic (deep crustal) lithologies. At the edge, SPAT is feldspathic in nature, likely due to the mixing of SPAT material with external FHT material(s) (e.g., Jolliff et al. 2000; Borst et al. 2012; Moriarty III and Pieters 2018).

Summary A number of distinct lithologies have been found on the lunar surface via remote investigations (Fig. 4), through in situ collection of material directly collected on the lunar surface, and through meteorite finds (and falls) on Earth. Beyond the compositional signatures described in the preceding sections, the lunar sample collection represents a variety of textures including crystalline samples, glasses (such as the glass beads sampled during Apollo missions), and breccias. Lunar rocks are also found as broken down, loose fragments in the regolith, in which particle sizes range from m-sized boulders to nm-sized dust. Evidence of impact events throughout the Moon’s history is recorded not only by basins and craters, but by the presence of breccias and impact melts throughout the lunar sample collections. Lithologies on the lunar surface which remain to be sampled directly include pure anorthosite which likely exists in the lunar crust (e.g., Ohtake et al. 2009), and dunite originating from the lunar mantle. At the time of writing, claims have been made regarding the potential identification of clasts in select meteorites which may represent these lithologies (see Korotev et al. 2009,

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Nagaoka et al. 2014, Treiman and Semprich 2019 for more information). Additional materials, yet to be explored on the lunar surface, may include ices which have been shown to be concentrated at the lunar poles (e.g., Anand et al. 2012). For additional discussion of specific lunar rock types, the reader is referred to Wang and Wu (2017), this volume.

References Anand M, Crawford IA, Balat-Pichelin M, Abanades S, van Westrenen W, Péraudeau G, Jaumann R, Seboldt W (2012) A brief review of chemical and mineralogical resources on the moon and their potential utilization. Planet Space Sci 74(1):42–48. https://doi.org/10.1016/ j.pss.2012.08.012 Borst AM, Foing BH, Davies GR, van Westrenen W (2012) Surface mineralogy and stratigraphy of the lunar south pole-Aitken basin determined from Clementine UV/VIS and NIR data. Planet Space Sci 68: 76–85. https://doi.org/10.1016/j.pss.2011.07.020 Elkins-Tanton LT, Burgess S, Yin Q-Z (2011) The lunar magma ocean: reconciling the solidification process with lunar petrology and geochronology. Earth Planet Sci Lett 304:326–336. https://doi.org/10.1016/j.epsl. 2011.02.004 Fortezzo CM, Spudis PD, Harrel SL (2020) Release of the digital unified global geologic map of the moon at 1: 5,000,000-Scale. Lunar Planet Sci. Conference LI abstract #2760 Garcia RF, Khan A, Drilleau M, Margerin L, Kawamura T, Sun D, Wieczorek MA, Rivoldini A, Nunn C, Weber RC, Marusiak AG, Longonné P, Nakamura Y, Zhu P (2019) Lunar seismology: an update on interior structure models. Space Sci Rev 215:50. https://doi.org/10. 1007/s11214-019-0613-y Geike A (1905) Geology of the moon. Nature 71:348–350. https://doi.org/10.1038/071348a0 Giguere TA, Taylor GJ, Hawke BR, Lucey PG (2000) The titanium contents of lunar mare basalts. Meteorit Planet Sci 35:193–200. https://doi.org/10.1111/j.1945-5100. 2000.tb01985.x Gillis JJ, Jolliff BL, Elphic RC (2003) A revised algorithm for calculating TiO2 from Clementine UVVIS data: a synthesis of rock, soil, and remotely sensed TiO2 concentrations. J Geophys Res 108(E2):1–18. https://doi. org/10.1029/2001JE001515 Gillis JJ, Jolliff BL, Korotev RL (2004) Lunar surface geochemistry: global concentrations of Th, K., and FeO as derived from lunar prospector and Clementine data. Geochim Cosmochim Acta 68(18):3791–3805. https://doi.org/10.1016/j.gca.2004.03.024 Hartmann WK (2019) History of the Termina cataclysm paradigm: epistemology of a planetary bombardment

917 that never (?) happened. Geosciences 9:285–363. https://doi.org/10.3390/geosciences9070285 Haskin LA, Gillis GJ, Korotev RL, Jolliff BL (2000) The materials of the lunar Procellarum KREEP terrane: a synthesis of data from geomorphological mapping, remote sensig, and sample analyses. J Geophys Res 105(E8):20403–20415. https://doi.org/10.1029/ 1999JE001128 Hawke BR, Peterson CA, Blewett DT, Bussey DBJ, Lucey PG, Taylor GJ, Spudis PD (2003) Distribution and modes of occurrence of lunar anorthosite. J Geophys Res 108(E6):5050. https://doi.org/10.1029/ 2002JE001890 Head JW (1976) Lunar volcanism in space and time. Rev Geophyscs Space Physics 14:265–300. https://doi.org/ 10.1029/RG014i002p00265 Head JW, Wilson L (2020) Rethinking Lunar Mare Basalt Regolith formation: new concepts of lava flow protolith and evolution of regolith thickness and internal structure. Geophys Res Lett 47:e2020GL088334. https:// doi.org/10.1029/2020GL088334 Hess PC, Parmentier EM (2001) Thermal evolution of a thicker KREEP liquid layer. J Geophys Res 106(E11):28023–28032. https://doi.org/10.1029/ 2000JE001416 Hiesinger H, Head JW III, Wolf U, Jaumann R, Neukum G (2003) Ages and stratigraphy of mare basalts in Oceanus Procellarum, Mare Nubium, Mare Cognitum, and Mare Insularum. J Geophys Res 108(E7):5065. https://doi.org/10.1029/2002JE001985 Hiesinger H, Head JW, Wolf U, Jaumann R, Neukum G (2011) Ages and stratigraphy of lunar mare basalts: a synthesis. In: Recent advances and current research issues in lunar stratigraphy, vol 477. Geological Society of America Special Paper, pp 1–51. https://doi.org/10. 1130/2011.2477(01) Hurwitz DM, Kring DA (2014) Differentiation of the south pole-Aitken basin impact melt sheet: implications for lunar exploration. J Geophys Res Lett 119:1110–1133. https://doi.org/10.1002/2013JE004530 Jolliff BL, Gillis JJ, Haskin LA, Korotev RL, Wieczorek MA (2000) Major lunar crustal terranes: surface expressions and crust-mantle origins. J Geophys Res 105(E2):4197–4216. https://doi.org/10.1029/ 1999JE001103 Kopal Z (1962) The internal constitution of the moon. Planet Space Sci 9(10):625–636. https://doi.org/10. 1016/0032-0633(62)90123-X Korotev RL, Zeigler RA, Jolliff BL, Irving AJ, Bunch TE (2009) Compositional and lithological diversity among brecciated lunar meteorites of intermediate iron concentration. Meteorit Planet Sci 44(9):1287–1322. https://doi.org/10.1111/j.1945-5100.2009.tb01223.x Lawrence DJ, Elphic RC, Feldman WC, Prettyman TH, Gasnault O, Maurice S (2002) Small-area thorium features on the lunar surface. J Geophys Res 108(E9). https://doi.org/10.1029/2003JE002050 Lemelin M, Lucey PG, Miljković K, Gaddis LR, Hare T, Ohtake M (2019) The compositions of the lunar crust

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Moon abundances estimated from UV/Vis reflectance. Icarus 296:216–238. https://doi.org/10.1016/j.icarus.2017. 06.013 Shearer CK, Hess PC, Wieczorek MA, Pritchard ME, Parmentier EM, Borg LE, Longhi J, Elkins-Tanton LT, Neal CR, Antonenko I, Canup RM, Halliday AN, Grove TL, Hager BH, Lee D-C, Wiechert U (2006) Thermal and magmatic evolution of the Moon. In: Jollif BL, Wieczorek MA, Shearer CK, Neal CR (eds) New views of the Moon. Reviews in Mineralogy & Geochemistry, 60(1), pp 365–518. https://doi.org/10.2138/ rmg.2006.60.4 Smith JV, Anderson AT, Newton RC, Olsen EJ, Wyllie PJ, Crewe AV, Isaacson MS, Johnson D (1970) Petrologic history of the moon inferred from petrography, mineralogy, and petrogenesis of Apollo 11 rocks. In: Proceedings of the Apollo 11 lunar science conference, pp 897–925 Smith DE, Zuber MT, Neumann GA, Lemoine FG, Mazarico E, Torrence MH, McGarry JF, Rowlands DD, Head JW III, Duxbury TH, Aharonsen O, Lucey PG, Robinson MS, Barnouin OS, Cavanaugh JF, Sun X, Liiva P, Mao, D.-d., Smith, J. C., Bartels, A. E. (2010) Initial observations from the Lunar Orbiter Laser Altimeter (LOLA). Geophys Res Lett 37: L18204. https://doi.org/10.1029/2010GL043751 Snape JF, Nemchin AA, Bellucci JJ, Whitehouse MJ, Tartèse R, Barnes JJ, Anand M, Crawford IA, Joy KH (2016) Lunar basalt chronology, mantle differentiation and implications for determining the age of the Moon. Earth Planet Sci Lett 451:149–158. https://doi.org/10. 1016/j.epsl.2016.07.026 Snape JF, Nemchin AA, Whitehouse MJ, Merle RE, Hopkinson T, Anand M (2019) The timing of basaltic volcanism at the Apollo landing sites. Geochim Cosmochim Acta 266:29–53. https://doi.org/10.1016/ j.gca.2019.07.042 Stöffler D, Ryder G, Ivanov BA, Artemieva NA, Cintala MJ, Grieve RAF (2006) Cratering history and lunar chronology. In: New views of the Moon: Reviews in mineralogy and geochemistry. Mineralogical Society of America, 60(1), pp 519–596. https://doi.org/10. 2138/rmg.2006.60.05 Tartèse R, Anand M, Gattacceca J, Joy K, Mortimer J, Pernet-Fisher JF, Russell S, Snape JF, Weiss BP (2019) Constraining the evolutionary history of the moon and the inner solar system: a case for new returned lunar samples. Space Sci Rev 2015:54. https://doi.org/10.1007/s11214-019-0622-x Taylor GJ (2009) Ancient lunar crust: origin, composition, and implications. Elements 5:17–22. https://doi.org/10. 2113/gselements.5.1.17 Treiman AH, Semprich J (2019) Dunite in lunar meteorite Northwest Africa 11421: petrology and origin. Lunar and planetary science conference L, abstract #1225 Vaughan WM, Head JW (2014) Impact melt differentiation in the south pole-Aitken basin: some observations and speculations. Planet Space Sci 91:101–106. https://doi. org/10.1016/j.pss.2013.11.010

Moon: Origin, Alternative Theories Wang X, Wu K (2017) Lunar Rocks. In: Cudnik B (ed) Encyclopedia of lunar science. Springer, Cham. https://doi.org/10.1007/978-3-319-05546-6_ 56-1 Warren PH, Wasson JT (1979) The Origin of KREEP. Rev Geophys Space Phys 17(1):73–88 Wieczorek MA (2009) The interior structure of the moon: what does geophysics have to say? Elements 5(1): 35–40. https://doi.org/10.2113/gselements.5.1.35 Wieczorek MA, Jolliff BL, Khan A, Pritchard ME, Weiss BP, Williams JG, Hood LL, Righter K, Neal CR, Shearer CK, McCallum IS, Tompkins S, Hawke BR, Peterson C, Gillis JJ, Bussey B (2006) The constitution and structure of the lunar interior. In: New views of the Moon: reviews in mineralogy & geochemistry. Mineralogical Society of America, 60(1), pp 221–364. https://doi.org/10.2138/rmg.2006.60.3 Wieczorek MA, Neumann GA, Nimmo F, Kiefer WS, Taylor JG, Melosh HJ, Phillips RJ, Solomon SC, Andrews-Hanna JC, Asmar SW, Knopliv AS, Lemoine FG, Smith DE, Watkins MM, Williams JG, Zuber MT (2013) The crust of the moon as seen by GRAIL. Science 339(6120):671–675. https://doi.org/10.1126/ science.1231530 Wood JA, Dickey JS, Marvin UB, Powell BN (1970) Lunar anorthosites and a geophysical model of the moon. Proceedings of the Apollo 11 Lunar Science Conference, pp 965–988 Zellner NEB (2017) Cataclysm no more: new views on the timing and delivery of lunar impactors. Orig Life Evol Biosph 47:261–280. https://doi.org/10.1007/ s11084-017-9536-3

Moon: Origin, Alternative Theories Caitlin Ahrens NASA Goddard Space Flight Center, Greenbelt, MD, USA

Definition The Moon’s origin and formation should be able to explain the current physical, dynamical, chemical, and internal thermal states. While the Giant Impactor Theory is the leading theory, we must also look into the earlier variations of lunar origins. This includes variants on three main possible scenarios: capture, rotational disruption, and co-accretion.

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Capture One particular idea was that the Moon was a strayed planetoid object, captured by Earth and circularized onto a stable orbit (Urey 1952; Mitler 1975). Most of these dynamical theories suggest a nebula that aids in the capture that has since dissipated (Gerstenkorn 1955). This nebula is likely to have vanished within ~1–10 Ma; this would not have been viable to physically capture, given the recognized current age of the Moon (Asphaug 2014). Asphaug (2014) states that instead a collision with Earth can decelerate a planetoid, to the point of capturing partial (or entirety) of it into orbit. A collision could have also led to the production of a swarm of captured planetesimals that would have eventually accreted into the Earth (Matsui and Abe 1986). Öpik (1972) suggested that a proto-moon approached Earth in a nearly parabolic orbit (defining a Keplerian orbit very similar to the Earth’s so that the Earth’s relative velocities are very low when the proto-moon is further away). If this proto-moon passed us with a perigee within the Roche distance, then it would have broken up, where the inner half would be captured and the other portion “rejected.” The captured pieces would then move in highly eccentric orbits and then be rapidly circularized by mutual collisions into a ring. Mild inelastic collisions would then eventually make these pieces coalesce (Mitler 1975). Mitler (1975) states that this mechanism for partial capture does not depend on any dissipative process, such as the configuration of the Earth-proto-moon or loss of orbital angular momentum. Mitler (1975) suggests that if the proto-moon were already differentiated (including an iron core), then it would be simple to shed the core, leaving material low in iron to reaccumulate into a Moon. Such advantages of having a capture-type mechanism for the origin of the Moon, as noted by Mitler (1975), were: (i) current hypotheses of planetary formation (Hallam and Marcus 1974); (ii) extensive cratering observed on the Moon, Mars, and Mercury; (iii) Age of the Moon being like that of Earth; and (iv) guarantees natural

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capture of material (or lack of certain materials, such as iron and siderophiles. However, Singer (1968) concludes that one of the major objections to the capture theory was namely the dissipation in a relatively short period of time considering great amounts of rotational kinetic energy inside the Earth and tidal disruption of the Moon.

Fission Darwin (1879) proposed that the Moon originated from a rapidly rotating Earth and can be modified by substitution of excessive rotation rate and unstable spin during the formation of the Earth’s core. The resulting instability would then “break” the Moon away. O’Keefe and Sullivan (1978) hypothesized that rapid differentiation changed Earth’s moment of inertia, resulting in an ejected debris ring of silicates. Singer (1968) objects that under the fission theory, the Moon would have formed initially in the equatorial plane of the Earth and that a multitude of moonlets would have formed within 10 Earth radii. From Wise (1963), the hypothesis of rotational fission during the Earth-core formation can be visualized as a very rapidly spinning Earth whereon a small addition to centrifugal force by increased spin rate would free a partial amount of the equator from its gravitational attraction. The spin rate to account for the rigidity and shape of the body should occur with a period of rotation of ~2.65 h (Jeans 1929). Assuming a proto-Earth of homogenous density, then settling, coalescing, compacting, or collapsing of the interior with time would reduce the moment of inertia (Wise 1963). Conservation of angular momentum to a newly differentiated Earth would then increase the rotational speed (Wise 1963). From this, a possible sequence of shapes can be taken (see Wise 1963). At low spin rates, the Earth’s shape starts as an oblate spheroid. With increasing spin rates, Jeans (1929) notes that the spheroid continues to flatten, whereupon unstable rotation begins. This transforms the oblate spheroid into an ellipsoid, revolving on its short axis. Finally with enough rotational acceleration, this

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shape would deform into a pear-shape, spinning with its long axis in the plane of rotation. The lunar embryo would be the “stem” oblate portion of the pear figure. This bulge would be composed primarily of Earth mantle material, with the outermost portion being primitive crustal components. Unstable rotation would cause the bulge to separate it from the proto-Earth. Both bodies would then reshape hydrostatically under their own gravitational forces to form the present Earth-Moon system (Wise 1963) (Fig. 1). Impact-induced fission is a variant of this theory (Ćuk and Stewart 2012). This includes a sequence of processes: (i) spin-up of Earth nearing disruption; (ii) fission of Earth’s mantle, triggered by smaller, energetic Theia-type

Moon: Origin, Alternative Theories, Fig. 1 Illustrative sequence of the lunar origins by fission during formation of the Earth’s core. (Adapted from Wise 1963)

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planetoids; and (iii) orbital evolution of the EarthMoon system to lose half of its angular momentum. From Wise (1963), the fission theory provided some major points of lunar origin (at least in the 1960s), such as (i) correctly predicting the direction of the Earth’s rotation being the same direction as the Moon’s; (ii) explain the relatively circular orbit of the Moon around the Earth; (iii) core-settling hypothesis predicts the observed specific gravity of the Moon; (iv) the same-facing position of the Moon due to the Moon accelerated out from the spinning proto-Earth; (v) explains the fundamental differences in the near versus far side of the Moon where the far side of the Moon would contain the remnants of lighter primitive Earth materials, whereas the nearside would be heavier mantle constituents; and (vi) the explanation of maximum moment of inertia along an axis pointing toward the Earth. Although this theory solves the isotropic mystery by essentially homogenizing the reservoirs, this hypothesis conflicts with iron segregation that occurred in the Moon’s early formation (Rudge et al. 2010). From Asphaug (2014), the Giant Impact Theory is considerably not that far removed from the concept of fission in that twoaccreting planet-like objects going from high to low inertia as their cores merge, transferring angular momentum to form a debris disk to coalesce.

Co-accretion This theory that Earth and the Moon co-accreted as a binary pair in the context of a solar nebula came about by Morishima and Watanabe (2001). In the plane of accretional and dynamical evolution of planet-satellite systems in a swarm of planetesimals on heliocentric orbits, a satellite with some initial value moves quickly toward the orbital radius, where accretion drag recompenses with tidal repulsion, and then grows toward equilibrium (Morishima and Watanabe 2001). Morishima and Watanabe (2001) proposed two co-accretion origins for the Moon. The first scenario describes the Moon starting as a small

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embryo (10
2 MM) and grew in a swarm of planetesimals with low velocity dispersion, so that large angular momentum was essentially supplied to planetary spinning. The rapid growth of the Earth (~106 years) was needed for the semimajor axis of the Moon to be kept so small that the Moon enlarged more rapidly under the benefit of “gravitational focusing” by the Earth. The second scenario describes that the Moon was formed by a giant impact occurring during Earth’s accretion. This impact supplied enough angular momentum as large as that of the present Earth-Moon system. This particular scenario is then a basis for the current (and favored) Giant-Impactor Theory. From Asphaug (2014), this co-accretion theory is not supported dynamically due to its dependence of a dynamic nebula and that W isotopes (and other chronometers; Touboul et al. 2007) have shown that the Moon must have formed at 4.5 Ga or later, long ahead of the proposed nebula dissipation. Co-formation, however, is somewhat consistent with the low-velocity collision that is required of the Giant-Impact Theory.

M References Asphaug E (2014) Impact origin of the Moon? Annu Rev Earth Planet Sci 42:551–578 Ćuk M, Stewart ST (2012) Making the Moon from a fastspinning Earth: a giant impact followed by resonant despinning. Science 338:1047–1052 Darwin GH (1879) A tidal theory of the evolution of satellites. Observatory 3:79–84 Gerstenkorn H (1955) Uber Gezeitenreibung beim Zweikorperproblem. Z Astrophys 36:245 Hallam M, Marcus AH (1974) Stochastic coalescence model for terrestrial planetary accretion. Icarus 21(1): 66–85 Jeans J (1929) Astronomy and cosmogony. Cambridge University Press, Cambridge, England Matsui T, Abe Y (1986) Origin of the Moon and its early thermal evolution. In: Origin of the Moon. Lunar and Planetary Institute, Houston, pp 453–468 Mitler HE (1975) Formation of an iron-poor moon by partial capture, or: yet another exotic theory of lunar origin. Icarus 24(2):256–268 Morishima R, Watanabe SI (2001) Two types of co-accretion scenarios for the origin of the Moon. Earth Planets Space 53(3):213–231 O’Keefe JA, Sullivan EC (1978) Fission origin of the Moon: cause and timing. Icarus 35:272–283

922 Öpik EJ (1972) Comments on lunar origin. Ir Astron J 10:190 Rudge JF, Kleine T, Bourdon B (2010) Broad bounds on Earth’s accretion and core formation constrained by geochemical models. Nat Geosci 3:439–443 Singer SF (1968) The origin of the Moon and geophysical consequences. Geophys J Int 15(1–2):205–226 Touboul M, Kleine T, Bourdon B, Palme H, Wieler R (2007) Late formation and prolonged differentiation of the Moon inferred from W isotopes in lunar metals. Nature 450:1206–1209 Urey HC (ed) (1952) The planets: their origin and development. Yale University Press, New Haven Wise DU (1963) An origin of the moon by rotational fission during formation of the earth's core. J Geophys Res 68(5):1547–1554

Moon: Seismicity Ceri Nunn Jet Propulsion Laboratory, California Institute of Technology, Pasadena, CA, USA

Definition The Moon is seismically active. During the years of operation of the Apollo seismometers, from 1969 to 1977, the mid-period seismometers recorded ~1750 meteoroid impacts, ~7400 deep moonquakes, and 28 shallow moonquakes, with many more events recorded on the other seismometers. The lack of atmosphere on the Moon means that meteoroids behave very differently from those impacting on Earth. Since there is no atmosphere to decelerate the meteoroids, they hit the Moon at full velocity and usually form spherically symmetrical craters. Deep moonquakes are small events that occur at depths of 700–1200 km and are likely driven by tides. Deep moonquakes often have similar waveforms, which can be assigned to groups by their similarity. All quakes within each group are thought to originate from a tight spatial cluster. Shallow moonquakes are rare events and often larger than other event types. Many authors locate shallow moonquakes to the upper mantle of the Moon, at depths between 50 and 200 km. Shallow moonquakes have some association with major basins and may be movement on

Moon: Seismicity

large faults surrounding the basins. Many thousands of thermal moonquakes, which are small events triggered by daily temperature changes, were also recorded.

Introduction Each Apollo mission to the Moon included a seismic experiment (▶ The Apollo Seismic Experiments). As Nakamura (2020) described, the seismometer at Apollo 11 recorded many seismic signals. However, the seismic signals looked distinct from signals seismologists were accustomed to seeing on Earth, often with unclear arrival phases and long durations. Until the seismologists saw the planned impact of the ascent stage of Apollo 12’s Lunar Module, recorded by Apollo 12’s seismometers, they were very unsure how to interpret the data. Nakamura further notes that the artificial seismic signal produced also looked very unfamiliar: The amplitude ramped up very gradually, and the event lasted for hours (Fig. 1). However, using this signal of known origin, location, and size, Latham et al. (1970b) began to assign the natural events from Apollo 11 and 12 to impacts or moonquakes. As seen in Fig. 1 and described by Latham and colleagues, the signals are long-duration, complex, emergent oscillations that lack the discrete phases of earthquake signals. In the early years of Apollo, it was somewhat surprising that the Moon was seismically active. Only meteoroid impacts were expected to be recorded, since there was a commonly held belief that the Moon was geologically dead (Urey 1952). In addition to thermal moonquakes, which tend to be small events detected locally by a seismometer, shallow and deep moonquakes were discovered. Figure 2 shows seismograms for a deep moonquake, a meteoroid impact, a shallow moonquake, and an artificial impact. Although the seismograms have much in common (like the artificial impacts, they are long-duration, complex, emergent oscillations which lack the discrete phases of earthquake signals), they also have several differences. Deep moonquakes are usually very small events with a sharper onset and faster

Moon: Seismicity

Moon: Seismicity, Fig. 1 The impact of the Apollo 12 Lunar Ascent Module after the crew left the Moon (recorded on the vertical component of the Apollo 12 midperiod seismometer). This signal helped to unlock the mystery of the lunar seismic signals and had many similarities with meteoroid impact signals already recorded at Apollo 11. The impact was approximately 73 km away

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(G. Latham et al. 1970a) and recorded at 22:17:17.7 on November 20, 1969 (Nunn et al. 2020 electronic supplement). Note that this is the time the signal was received on Earth, and therefore the impact was approximately 1.3 s earlier. Lognonné et al. (2003) estimated the arrival of the first seismic signal to be 24.7  10 s after the impact time

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Moon: Seismicity, Fig. 2 Examples of a deep moonquake, a meteoroid impact, a shallow moonquake, and an artificial impact event. The events were recorded at Apollo 12 on three components (MHZ, MH1, and MH2). Timing is relative to the first arrival, which is indicated on each of

the events. The y-axis scale is in raw digital units (DU) used by the seismometer, and the scale is different for each of the events. On the highest amplitude signal (the artificial impact), the signal was clipped. (Reproduced from Nunn et al. (2020))

rise time. Meteoroid impacts have much in common with artificial impacts since both are caused by an impactor that hits the Moon at high velocity. Shallow moonquakes have a slower rise time as well as significant high-frequency energy. Nakamura et al. (1981, and updates) categorized each event and made a catalog of lunar events (see also Nunn et al. (2020) for the current catalog). All types of moonquakes have significant uncertainties associated with the arrival times of even the P and S phases (in many cases, even these cannot be determined at all). Additionally, the network was small and covered only part of the

nearside. Consequently, the estimates of moonquake locations are highly uncertain (Garcia et al. 2019).

Artificial Impacts To provide large seismic sources to probe the lunar interior, the flight engineers deliberately crashed the Saturn V’s third stage boosters (S-IVB) and the ascent stages of the Lunar Module into the Moon. These nine events are valuable since the location, the timing of the impact, and

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Moon: Seismicity

Moon: Seismicity, Fig. 3 The 30 m crater left by the third stage of the Apollo 14 Saturn V, which was intentionally crashed into the Moon to serve as an energy source to probe the Moon’s interior, photographed by the Lunar

Reconnaissance Orbiter Camera (LROC 2013). The Artificial Impact in Fig. 2 shows the seismograms produced by the event. (Photo Credit: NASA/GSFC/Arizona State University)

the impact energy are known. The one exception is Apollo 16’s S-IVB impact, where the timing and location were unknown because the tracking was lost. However, the impact crater was later found (Plescia et al. 2016). Seismologists use these events to calibrate other signals and to refine models of the lunar interior. Figure 3 shows the 30 m crater left by Apollo 14’s S-IVB, and the Artificial Impact in Fig. 2 shows the seismogram resulting from the impact. Using images from the Lunar Reconnaissance Orbiter Camera, Plescia et al. (2016), Wagner et al. (2017), and Stooke (2017) estimated the precise locations of many of the impacts.

originating from either comets or asteroids, and estimated the mass for the meteoroids to range from 100 g to 100 kg. The waveforms of meteoroid and artificial impacts differ significantly from fault-generated quakes. They do not have a double-couple source. Since the Moon has no significant atmosphere, impacts have high velocities, and the impactor tends to fragment and vaporize. Teanby and Wookey (2011) noted that this leads to the creation of radially symmetric craters, except for very low-angle impacts (with respect to the horizontal). Since the impacts are surface events, seismic waves propagate through the regolith and megaregolith twice, once in the source region and once beneath the seismic station. The travel path of the seismic wave results in different scattering features and generates more gradual signal onset and longer-lasting signals than shallow or deep moonquakes. While some experiments have studied the seismic characteristics of impacts (e.g., McGarr et al. (1969) and Yasui et al. (2015)), observations from Apollo are still the only example of impacts on a body without an atmosphere and provide a unique opportunity to investigate the source mechanism.

Meteoroid Impacts About 4000 objects of more than 1 kg impact the Moon per year (Lognonne et al. 2009). With the earliest Apollo data, Latham et al. (1970b) found that many of the recorded seismic events were meteoroid impacts. Nakamura et al. (1981) attributed more than 1700 events recorded during the operation of the Apollo stations to meteoroid impacts. Oberst and Nakamura (1991), Lognonné et al. (2009), and Oberst et al. (2012) used seismically detected lunar impacts to estimate the flux of meteoroids in the Earth-Moon system. Those estimates were comparable to those obtained from other means (Daubar et al. 2018). Oberst and Nakamura (1991) found two distinct classes of meteoroids impacting the Moon,

Shallow Moonquakes Nakamura et al. (1981) designated 28 events as shallow moonquakes in their seismic event catalog. Therefore, shallow moonquakes are the rarest

Moon: Seismicity

event category. They often have larger magnitudes than the other naturally occurring events. Oberst (1987) estimated the equivalent bodywave magnitudes to be between 3.6 and 5.8, and he estimated mechanisms with unusually high stress drops. They are clearly visible on the short-period seismograms and contain highfrequency seismic energy (which drops off around 8 Hz at the high-frequency limit of the instrument). Nakamura (1977) showed no correlation between shallow moonquakes and the tides acting on the Moon. Instead, there are many similarities between these quakes and intraplate earthquakes on Earth (Nakamura 1980). There is some disagreement about the depth ranges for the shallow events, although most authors agree that they originate from the upper mantle. Using the amplitude decay function of the shallow moonquakes, Nakamura et al. (1979) argued that they are shallower than 200 km but deeper than the crust-mantle boundary. Khan et al. (2000) also preferred an upper mantle origin (at depths between 50 and 220 km). Garcia et al.’s (2011) VPREMOON model locates the quakes from depths of 0 to 168 km. Gillet et al. (2017) used the scattering properties of these quakes to suggest that they originate from a depth of about 50 km  20 km. Nakamura et al. (1979) found a correlation between their locations and the lunar basins. They suggested that the shallow moonquakes may be distributed along circular faults around impact basins, a conclusion also supported by Gillet et al. (2017).

Deep Moonquakes Reflecting on the initial analysis of the Apollo seismic data, Nakamura (2020) noted that the seismologists were surprised to find that the sources of the seismic events with distinct P- and S-wave arrivals recorded by the seismometers were not anywhere near the surface of the Moon as we might expect for seismicity related to tectonic processes. Instead, the quakes were at great depths, deeper than deep earthquakes, and about halfway to the center of the Moon. These quakes,

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which became classified as deep moonquakes, are the most numerous and occur at depths of 700 to 1200 km (Nakamura 2005; Nakamura et al. 1982). Different quakes in a given group of moonquakes have similar waveforms, which led to the theory that they originate from tightly clustered source regions (or “nests”). Several authors classified the moonquakes into numbered clusters (e.g., Bulow et al. 2007; Lognonné et al. 2003; Nakamura 1978). Nakamura (2005) identified at least 165 different source regions, mainly on the nearside of the Moon. The largest cluster, A1, contains over 400 quakes. The A1 group is large enough to distinguish events with slightly different waveform subgroups (Gagnepain-Beyneix et al. 2006), and the delays between P and S arrival times obtained by correlation implied that the distance between sources was less than 1 km. Using single-link cluster analysis, Nakamura (2003) correlated every pair of events. He found that waveforms from events that belonged to one source region were highly correlated, while those belonging to different source regions correlated to a lower degree. However, he was surprised to find that events thought to belong to separate source regions still correlated. Early on, Lammlein et al. (1974) noted an association between the times of deep moonquakes and the tidal phases of the Moon. Analysis of individual clusters by Frohlich and Nakamura (2009) shows tidal periodicity for each cluster, but not necessarily the same dependence on the tidal cycle for all clusters. Although there is a consensus that deep moonquakes are tidally triggered, the precise cause remains unclear. Possible mechanisms include the presence of fluids (especially water) in the source region (Saal et al. 2008) or partial melts (Frohlich and Nakamura 2009). Calculated tidal stresses are strongest from 600 to 1200 km (e.g., Cheng and Toksöz (1978)), where the deep moonquakes occur. Although most deep moonquake nests are located on the nearside of the Moon, Nakamura (2005) located 30 nests to the farside. He found that none of the farside nests were within about 40 of the antipode of the Moon. He suggested that the deep interior of the Moon

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Moon: Seismicity, Fig. 4 Thermal moonquake examples, detected by Apollo 17’s Lunar Seismic Profiling Experiment, falling into the categories impulsive,

intermediate, and emergent. (Reproduced under Creative Commons License from Dimech et al. (2017))

severely attenuates or deflects seismic waves, accounting for the apparent absence of deep moonquakes from the antipode.

References

Thermal Moonquakes Duennebier and Sutton (1974) showed that most of the many thousands of seismic events recorded on the short-period seismometers were small local moonquakes triggered by daily temperature changes, and classified these events as thermal moonquakes. Figure 4 shows examples of thermal moonquakes detected by Apollo 17’s Lunar Seismic Profiling Experiment. Dimech et al. (2017) made nearly 50,000 detections of these small quakes. Of these, 5% were impulsive, 27% were intermediate, and the remaining 69% were emergent. The events occurred periodically, with a sharp double peak at sunrise and a broad single peak at sunset. Weber et al. (2017) suggested that the difference in the seismogram shape reflected their relative distance from the array, with impulsive events being closest and emergent events being further away. The impulsive events occur primarily at sunrise, possibly representing the thermal “pinging” of the nearby lunar lander. In contrast, emergent events occur at sunset, perhaps representing cracking or slumping in more distant surface rocks and regolith. Acknowledgments The work was carried out at the Jet Propulsion Laboratory, California Institute of Technology, under a contract with the National Aeronautics and Space Administration (80NM0018D0004).

Bulow RC, Johnson CL, Bills BG, Shearer PM (2007) Temporal and spatial properties of some deep moonquake clusters. J Geophys Res 112. https://doi. org/10.1029/2006JE002847 Cheng CH, Toksöz MN (1978) Tidal stresses in the Moon. J Geophys Res 83:845. https://doi.org/10.1029/ JB083iB02p00845 Daubar I, Lognonné P, Teanby NA, Miljkovic K, Stevanović J, Vaubaillon J, Kenda B, Kawamura T, Clinton J, Lucas A, Drilleau M, Yana C, Collins GS, Banfield D, Golombek M, Kedar S, Schmerr N, Garcia R, Rodriguez S, Gudkova T, May S, Banks M, Maki J, Sansom E, Karakostas F, Panning M, Fuji N, Wookey J, van Driel M, Lemmon M, Ansan V, Böse M, Stähler S, Kanamori H, Richardson J, Smrekar S, Banerdt WB (2018) Impact-seismic investigations of the InSight mission. Space Sci Rev 214. https://doi.org/ 10.1007/s11214-018-0562-x Dimech J-L, Knapmeyer-Endrun B, Phillips D, Weber RC (2017) Preliminary analysis of newly recovered Apollo 17 seismic data. Results Phys 7:4457–4458. https://doi. org/10.1016/j.rinp.2017.11.029 Duennebier F, Sutton GH (1974) Thermal moonquakes. J Geophys Res 79:4351–4363. https://doi.org/10.1029/ JB079i029p04351 Frohlich C, Nakamura Y (2009) The physical mechanisms of deep moonquakes and intermediate-depth earthquakes: how similar and how different? Phys Earth Planet Interiors 173:365–374. https://doi.org/10.1016/ j.pepi.2009.02.004 Gagnepain-Beyneix J, Lognonné P, Chenet H, Lombardi D, Spohn T (2006) A seismic model of the lunar mantle and constraints on temperature and mineralogy. Phys Earth Planet Interiors 159:140–166. https://doi.org/10.1016/j.pepi.2006.05.009 Garcia RF, Gagnepain-Beyneix J, Chevrot S, Lognonné P (2011) Very preliminary reference Moon model. Phys Earth Planet Interiors 188:96–113. https://doi.org/10. 1016/j.pepi.2011.06.015 Garcia RF, Khan A, Drilleau M, Margerin L, Kawamura T, Sun D, Wieczorek MA, Rivoldini A, Nunn C, Weber RC, Marusiak AG, Lognonné P, Nakamura Y, Zhu

Moon: Seismicity P (2019) Lunar seismology: an update on interior structure models. Space Sci Rev 215. https://doi.org/10. 1007/s11214-019-0613-y Gillet K, Margerin L, Calvet M, Monnereau M (2017) Scattering attenuation profile of the Moon: implications for shallow moonquakes and the structure of the megaregolith. Phys Earth Planet Interiors 262:28–40. https:// doi.org/10.1016/j.pepi.2016.11.001 Khan A, Mosegaard K, Rasmussen KL (2000) A new seismic velocity model for the Moon from a Monte Carlo inversion of the Apollo lunar seismic data. Geophys Res Lett 27:1591–1594. https://doi.org/10.1029/ 1999GL008452 Lammlein DR, Latham GV, Dorman J, Nakamura Y, Ewing M (1974) Lunar seismicity, structure, and tectonics. Rev Geophys 12:1–21. https://doi.org/10.1029/ RG012i001p00001 Latham G, Ewing M, Dorman J, Press F, Toksoz N, Sutton G, Meissner R, Duennebier F, Nakamura Y, Kovach R, Yates M (1970a) Seismic data from manmade impacts on the Moon. Science 170:620–626. https://doi.org/10.1126/science.170.3958.620 Latham GV, Ewing M, Press F, Sutton G, Dorman J, Nakamura Y, Toksoz MN, Wiggins R, Derr J, Duennebier F (1970b) Passive Seismic Experiment. Science 167:455–457. https://doi.org/10.1126/science. 167.3918.455 Lognonné P, Gagnepain-Beyneix J, Chenet H (2003) A new seismic model of the Moon: implications for structure, thermal evolution and formation of the Moon. Earth Planet Sci Lett 211:27–44. https://doi. org/10.1016/S0012-821X(03)00172-9 Lognonne P, Le Feuvre M, Johnson CL, Weber RC (2009) Moon meteoritic seismic hum: steady state prediction. J Geophys Res-Planets 114:E12003. https:// doi.org/10.1029/2008JE003294 LROC (2013) Apollo 14 S-IVB, http://www.lroc.asu.edu/ featured_sites/view_site/25 McGarr A, Latham GV, Gault DE (1969) Meteoroid impacts as sources of seismicity on the Moon. J Geophys Res 74:5981–5994. https://doi.org/10. 1029/JB074i025p05981 Nakamura Y (1977) HFT events: shallow moonquakes? Phys Earth Planet Interior 14:217–223. https://doi.org/ 10.1016/0031-9201(77)90174-1 Nakamura Y (1978) A1 moonquakes – source distribution and mechanism. In: Proc Lunar Planet Sci Conf 9th, pp. 3589–3607. https://articles.adsabs.harvard.edu/pdf/ 1978LPSC....9.3589N Nakamura Y (1980) Shallow moonquakes – how they compare with earthquakes. In: Proc Lunar Planet Sci Conf 11th, pp. 1847–1853. https://articles.adsabs. harvard.edu/pdf/1980LPSC...11.1847N Nakamura Y (2003) New identification of deep moonquakes in the Apollo lunar seismic data. Phys Earth Planet Interior 139:197–205. https://doi.org/10.1016/j. pepi.2003.07.017

927 Nakamura Y (2005) Farside deep moonquakes and deep interior of the Moon. J Geophys Res 110. https://doi. org/10.1029/2004JE002332 Nakamura Y (2020) Rebirth of extraterrestrial seismology. Nat Geosci 13:178–179. https://doi.org/10.1038/ s41561-020-0551-z Nakamura Y, Latham GV, Dorman HJ, Ibrahim A-BK, Koyama J, Horvath P (1979) Shallow moonquakes – depth, distribution and implications as to the present state of the lunar interior. In: Proc lunar and planetary science conference 10th lunar and planetary science conference proceedings, vol 10, pp 2299–2309. https://articles.adsabs.harvard.edu/pdf/1979LPSC...10. 2299N Nakamura Y, Latham GV, Dorman HJ, Harris JE (1981) Passive seismic experiment long-period event catalog. Final version. UTIG technical report no. 18. Galveston Geophysics Laboratory of the University of Texas at Austin, Marine Science Institute Nakamura Y, Latham GV, Dorman HJ (1982) Apollo Lunar seismic experiment – final summary. J Geophys Res Solid Earth 87:A117–A123. https://doi.org/10. 1029/JB087iS01p0A117 Nunn C, Garcia RF, Nakamura Y, Marusiak AG, Kawamura T, Sun D, Margerin L, Weber R, Drilleau M, Wieczorek MA, Khan A, Rivoldini A, Lognonné P, Zhu P (2020) Lunar seismology: a data and instrumentation review. Space Sci Rev 216:89. https://doi.org/10.1007/s11214-020-00709-3 Oberst J (1987) Unusually high stress drops associated with shallow moonquakes. J Geophys Res 92: 1397–1405. https://doi.org/10.1029/ JB092iB02p01397 Oberst J, Nakamura Y (1991) A search for clustering among the meteoroid impacts detected by the Apollo lunar seismic network. Icarus 91:315–325. https://doi. org/10.1016/0019-1035(91)90027-Q Oberst J, Christou A, Suggs R, Moser D, Daubar IJ, McEwen AS, Burchell M, Kawamura T, Hiesinger H, Wünnemann K, Wagner R, Robinson MS (2012) The present-day flux of large meteoroids on the lunar surface – a synthesis of models and observational techniques. Planet Space Sci 74:179–193. https://doi.org/ 10.1016/j.pss.2012.10.005 Plescia JB, Robinson MS, Wagner R, Baldridge R (2016) Ranger and Apollo S-IVB spacecraft impact craters. Planet Space Sci 124:15–35. https://doi.org/10.1016/j. pss.2016.01.002 Saal AE, Hauri EH, Cascio ML, Van Orman JA, Rutherford MC, Cooper RF (2008) Volatile content of lunar volcanic glasses and the presence of water in the Moon’s interior. Nature 454:192–195. https://doi.org/ 10.1038/nature07047 Stooke PJ (2017) Spacecraft impacts on the Moon: Change 1’, Apollo LM ascent stages. In: 48th Lunar Planet. Sci. Conf., Abstract #1031. https://www.hou.usra.edu/meet ings/lpsc2017/pdf/1031.pdf

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928 Teanby NA, Wookey J (2011) Seismic detection of meteorite impacts on Mars. Phys Earth Planet Interiors 186: 70–80. https://doi.org/10.1016/j.pepi.2011.03.004 Urey H (1952) The planets: their origin and development. Yale University Press, New Haven Wagner RV, Nelson DM, Plescia JB, Robinson MS, Speyerer EJ, Mazarico E (2017) Coordinates of anthropogenic features on the Moon. Icarus. Lunar Reconnaissance Orbiter – Part II 283:92–103. https://doi.org/ 10.1016/j.icarus.2016.05.011 Weber RC, Dimech JL, Phillips D, Molaro J, Schmerr NC (2017) A new moonquake catalog from Apollo 17 seismic data I: Lunar seismic profiling experiment: thermal

Multi-ring Basin moonquakes and implications for surface processes. In: AGU Fall Meeting Abstracts, p. P44B-09 Yasui M, Matsumoto E, Arakawa M (2015) Experimental study on impact-induced seismic wave propagation through granular materials. Icarus 260:320–331. https://doi.org/10.1016/j.icarus.2015.07.032

Multi-ring Basin ▶ Impact Processes on the Moon

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Noble Gases Vincent Eke and Jacob Kegerreis Institute for Computational Cosmology, Durham University, Durham, UK

Definition Noble gases are elements in group 18 of the periodic table. This article will consider the noble gases present in the lunar exosphere and their interaction with the lunar surface.

Introduction Noble gas elements have full outer electron shells, making them inert. Consequently, observations of noble gases in the lunar exosphere are more readily related to their sources and sinks than is the case for elements that are more chemically active. The study of the lunar noble gas exosphere thus informs us about the solar wind, the lunar interior and outgassing events, the interactions between atoms and the lunar regolith, and the sequestration of volatiles in polar cold traps (Stern 1999; Wieler and Heber 2003). To date, helium (He), neon (Ne), and argon (Ar) have all been directly detected in the lunar exosphere. Helium and argon were first identified in the lunar exosphere by the mass spectrometer in the Lunar Atmosphere Composition Experiment © Springer Nature Switzerland AG 2023 B. Cudnik (ed.), Encyclopedia of Lunar Science, https://doi.org/10.1007/978-3-319-14541-9

(LACE, Hoffman et al. 1973). It was also suggested that a peak at mass 20 was caused by neon, but this was later deemed to be plausibly due to outgassing of H218O (Hodges et al. 1973; Stern 1999). Due to outgassing from hardware placed in the vicinity of the LACE, the detector was typically switched off shortly after sunrise each morning, so the results only permitted studies of the nighttime exosphere. In contrast, the neutral mass spectrometer (NMS) on board NASA’s Lunar Atmosphere and Dust Environment Explorer (LADEE, Elphic et al. 2014) not only provided the discovery of neon in the lunar exosphere, but also enabled studies of the variation of the helium, neon, and argon exospheres during the lunar day thanks to its low-altitude retrograde orbit (Benna et al. 2015). The presence of a fourth noble gas, radon (Rn), has been inferred via measurements of α particles with energies characteristic of the radioactive isotope 222Rn and its daughter polonium isotopes 218 Po, 214Po, and 210Po. These isotopes are part of the uranium series decay chain taking 238U to 206 Pb, and radon is of particular interest because it is gaseous and can escape into the exosphere. The first orbital detection of 210Po was made by Gorenstein and Bjorkholm (1972) with the α Particle Spectrometer orbital experiment on Apollo 15. Additional measurements from Apollo 16 showed that there was spatial variation in the distributions of both 210Po and 222Rn (Gorenstein et al. 1974). Over 30 years later, the Lunar Prospector α Particle Spectrometer “marginally

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detected” 218Po, as well as yielding the first global maps of the 222Rn and 210Po distributions (Lawson et al. 2005). While the sources and sinks of the various noble gases may differ, the atoms have in common the mechanism by which they evolve through the lunar exosphere, namely, a series of “hops” involving thermal accommodation to the surface followed by a randomly directed emission into a ballistic trajectory (Chamberlain 1963). The size of the hop depends upon both the local temperature of the regolith and the mass of the atom. Also, while helium and neon are noncondensable gases, the heavier argon and radon atoms do tend to adsorb to the colder overnight surface as a result of their stronger van der Waals forces. This leads to qualitatively very different diurnal variability of the exospheres of these condensable gases. The remainder of this chapter will focus on each of the four noble gases known to be present in the lunar exosphere, describing in turn what has been learned from these elements.

Helium The main source of helium for the lunar exosphere is α particles in the solar wind. A second nonnegligible source is outgassing of helium from the Moon, produced by the uranium and thorium decay chains. As a result of its low mass, when a helium atom thermally accommodates to the hot, daytime lunar surface, it can acquire sufficient energy to become gravitationally unbound. This thermal escape is the main loss mechanism of helium from the lunar exosphere, with photoionization playing a smaller role. The typical lifetime of exospheric helium atoms from source to sink is
5 days, this being the decay timescale of the helium abundance measured by the Lunar Reconnaissance Orbiter (LRO) Lyman Alpha Mapping Project (LAMP) ultraviolet spectrograph when the Moon passed through the Earth’s magnetotail and the solar wind source was effectively switched off (Feldman et al. 2012). Typical helium atom hops traverse
1000 km in
15 min, so these atoms quickly lose any

Noble Gases

memory of the location at which they were injected into the exosphere. Despite the subsolar point receiving the highest solar wind He2+ flux, the relatively high temperatures and large resulting hops preferentially diffuse helium atoms away from this point to colder parts of the surface. As a consequence, the inferred exospheric helium density at the surface peaks just over an hour after midnight at 6.5  104 atoms cm3, a value almost 30 times that seen at noon (Hoffman et al. 1973; Benna et al. 2015). Lower regolith temperatures occur later in the night, after the peak in the helium exospheric density, but these locations are sufficiently close to the sunrise terminator that they receive a relative lack of atoms hopping back into the shade (Hurley et al. 2016). Contemporaneous measurements of the solar wind α flux and the exospheric helium density show that, once integrated over an assumed 4.5day thermal escape timescale, temporal variations in the recent solar wind source correlate well with the instantaneous exospheric helium density. In this way, using data from the Acceleration, Reconnection, Turbulence, and Electrodynamics of Moon’s Interaction with the Sun (ARTEMIS) twin spacecraft and the LADEE NMS, Benna et al. (2015) inferred that the solar wind contributed 85% of the helium exosphere, with a constant outgassing component providing the rest. This comparison relies upon knowing what fraction of incident solar wind α particles are retained in the exosphere rather than being backscattered out of the system. Hurley et al. (2016) and Grava et al. (2016) included LRO/LAMP measurements into their exospheric models to constrain this retained fraction to 50–60%. The LRO/LAMP measurements extend our understanding of the lunar exosphere beyond the near-equatorial zone probed by the LADEE NMS.

Neon The neon in the lunar exosphere is delivered by the solar wind, sharing an isotope ratio of 20 Ne/22Ne ¼ 14 (Benna et al. 2015). Neon atoms

Noble Gases

are too massive for thermal escape to be a significant loss mechanism, so their lifetimes in the exosphere are set by the ~6-month photoionization timescale. This means short-timescale variability in the solar wind is barely reflected in the neon exospheric density. Neon, like helium, is a noncondensable gas so the diurnal variation of density is qualitatively similar to that of helium. Typical hop sizes and corresponding flight times for the random walks carried out by the atoms through the exosphere pffiffiffiffi scale like 1/m and 1= m , where m is the atomic mass. Thus, the main difference between the neon and helium diurnal variation is that the factor of 5 shorter hops leads to a peak concentration of neon that occurs later in the night. The LADEE NMS measured this maximum to be ~3  104 cm3, similar to the helium peak density (Benna et al. 2015). This similarity results from the combination of a lower solar wind neon flux and a longer neon lifetime in the exosphere. Benna et al. (2015) presented the LADEE NMS results for just the nighttime neon density, because of the difficulty in removing the 40Ar2+ contamination from the 20Ne+ signal during the day.

Argon About 10% of the argon in the lunar exosphere is 36 Ar, originating in the solar wind. The remainder is 40Ar, which is the result of radioactive decay of 40 K and subsequent outgassing from the lunar interior (Hoffman et al. 1973). All three of charge exchange with solar protons, photoionization, and cold-trapping represent significant loss mechanisms for argon (Grava et al. 2015). Uncertainties in the details of cold-trapping produce a large uncertainty in the typical lifetime of argon atoms in the lunar exosphere, but it is likely to be at least a few weeks and not more than a few months (Grava et al. 2015; Kegerreis et al. 2017). Unlike helium and neon, argon is condensable on the cold, nighttime surface. This reduces the exospheric density overnight to an extent where the LADEE NMS data become difficult to interpret away from the terminators (Kegerreis et al.

931

2017). Fortunately, this is the time of day when LACE observations were not compromised by outgassing from hardware (Hodges 1975), so data are available covering all local times at equatorial latitudes. Peak exospheric argon densities are found at both terminators as a result of the diffusion of atoms from the hotter dayside leading to concentrations just after sunset and sunrise. The sunrise peak is the larger one, where many adsorbed atoms suddenly become activated by the increase in temperature, producing densities up to 105 cm3 (Benna et al. 2015). In models of the argon exosphere, the peak location in local time is very sensitive to the desorption energy required to activate the argon atoms (Kegerreis et al. 2017). That the measured peak location is insensitive to selenographic longitude implies that desorption energies do not vary significantly around the equatorial region. The amplitude of the sunrise peak does depend on selenographic longitude, as does the argon exospheric density at other times of day, albeit to a lesser extent. An “argon bulge” is seen in the western maria that increases the sunrise density by a factor of 4 (Benna et al. 2015). As the atoms diffuse away from this region after it experiences sunrise, an excess is observed to sweep through neighboring longitudes at nearby local times. This behavior can only be well reproduced if there is a localized source that is responsible for recharging the bulge (Kegerreis et al. 2017). The LADEE NMS argon data also show variability of
40% over the course of the 128-day measurement period. One elegant possible explanation is that this reflects a seasonal variation in the cold-trap surface area, albeit only observed for just under half a period (Hodges and Mahaffy 2016). While this fits well with the long-term variation detected by the LADEE NMS, the same explanation does not fit the LACE data, which showed a drop by a factor of 2 over a 4-month period (Hodges 1975; Kegerreis et al. 2017). However, Hodges and Mahaffy (2016) suggested that the decrease measured by LACE could have been the result of changing detector sensitivity. Time-variability of the argon source could also provide part of the explanation, with

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moonquakes (Grava et al. 2015) and diffusion (Killen 2002) among possible candidates for the outgassing mechanism.

Noble Gases

supports the inference of sporadic outgassing events being responsible.

Summary Radon Radon is created by radioactive decay of radium within the Moon. In particular, 222Rn is part of the decay chain of 238U. It has a half-life of only
3.8 days, so must be delivered rapidly into the lunar exosphere if it is to be observed before it decays to 218Po. The production of α particles from the decay of the daughter isotope 210Po is delayed by the
21-year half-life of 210Pb, which is an intermediate link in the decay chain (Lawson et al. 2005). 222Rn, on account of its high mass, does not diffuse very far over the lunar surface before decaying. Therefore, the spatial distribution of 210Po decays can be used in conjunction with that of 222Rn to probe both temporal and spatial variability of 222Rn outgassing. Locations at which 222Rn and 210Po have been observed to be decaying are not uniformly distributed across the lunar surface. Apollo 15 and Apollo 16 measurements showed excess 222Rn over Aristarchus crater, whereas 210Po was more abundant over Grimaldi crater and at the edges of most of the maria crossed by the spacecraft groundtracks (Gorenstein et al. 1974). Lunar Prospector similarly found excess 222Rn above Aristarchus, with a new detection over Kepler crater. However, there was no longer a significant excess of 210Po above Grimaldi and only a few maria margins showed an enhancement, including Serenitatis, which was the only one for which Apollo data did not show an excess (Lawson et al. 2005). The depths at which the liberated 222Rn is created and the manner by which it reaches the surface remain uncertain, although both the observations and laboratory studies suggest that sporadic venting appears more likely than a more continuous diffusion process (Friesen and Adams 1976). Furthermore, Aristarchus is the location most frequently associated with lunar transient phenomena (Buratti et al. 2000), so this

The helium and neon exospheres predominantly inform us about the interaction of the solar wind with the lunar surface. The outgassed helium component has a source rate that is known to within a factor of
2, but its spatial distribution and what that might imply about the lunar interior remains poorly quantified. The study of the argon and radon components of the lunar exosphere provides constraints on outgassing mechanisms that in turn should yield information about the lunar interior. While the behavior of these elements once in the exosphere is reasonably well understood, much remains to be learned about their injection into the system. The main source of uncertainty in the exospheric evolution of all noble gases is the extent of the cold-trapping of argon at the poles. Sending a rover to prospect in these particularly cold permanently shaded regions would be an effective way to investigate this in more detail. Longer term monitoring of the argon exospheric density could test if the argon exospheric density really does vary because of seasonally varying cold traps.

Cross-References ▶ Contribution of Surface Processes to the Lunar Exosphere ▶ Lunar Atmosphere ▶ Lunar Atmosphere, Composition ▶ Lunar Atmosphere, Source and Loss Processes ▶ Lunar Atmosphere, Transport and Storage of Volatiles

References Benna M, Mahaffy PR, Halekas JS, Elphic RC, Delory GT (2015) Variability of helium, neon, and argon in the lunar exosphere as observed by the LADEE NMS instrument. Geophys Res Lett 42:3723–3729

Non-mare Volcanism Buratti BJ, McConnochie TH, Calkins SB, Hillier JK, Herkenhoff KE (2000) Lunar transient phenomena: what do the clementine images reveal? Icarus 146:98–117 Chamberlain JW (1963) Planetary coronae and atmospheric evaporation. Planet Space Sci 11:901–960 Elphic RC, Delory GT, Hine BP, Mahaffy PR, Horanyi M, Colaprete A, Benna M, Noble SK (2014) The lunar atmosphere and dust environment explorer mission. Space Sci Rev 185:3–25 Feldman PD, Hurley DM, Retherford KD, Gladstone GR, Stern SA, Pryor W, Parker JW, Kaufmann DE, Davis MW, Versteeg MH, LAMP Team (2012) Temporal variability of lunar exospheric helium during January 2012 from LRO/LAMP. Icarus 221:854–858 Friesen LJ, Adams JAS (1976) Low pressure radon diffusion – a laboratory study and its implications for lunar venting. Geochim Cosmochim Acta 40:375–380 Gorenstein P, Bjorkholm P (1972) Observation of lunar radon emanation with the Apollo 15 alpha particle spectrometer. In: Metzger AE, Trombka JI, Peterson LE, Reedy RC, Arnold JR (eds) Lunar and planetary science conference proceedings, Houston, vol 3, p 2179 Gorenstein P, Golub L, Bjorkholm P (1974) Radon emanation from the moon, spatial and temporal variability. Moon 9:129–140 Grava C, Chaufray J-Y, Retherford KD, Gladstone GR, Greathouse TK, Hurley DM, Hodges RR, Bayless AJ, Cook JC, Stern SA (2015) Lunar exospheric argon modeling. Icarus 255:135–147 Grava C, Retherford KD, Hurley DM, Feldman PD, Gladstone GR, Greathouse TK, Cook JC, Stern SA, Pryor WR, Halekas JS, Kaufmann DE (2016) Lunar exospheric helium observations of LRO/LAMP coordinated with ARTEMIS. Icarus 273:36–44 Hodges RR Jr, Hoffman JH, Johnson FS, Evans DE (1973) Composition and dynamics of lunar atmosphere. In: Lunar and planetary science conference proceedings, Houston, vol 4, p 2855

933 Hodges RR Jr (1975) Formation of the lunar atmosphere. Moon 14:139–157 Hodges RR, Mahaffy PR (2016) Synodic and semiannual oscillations of argon-40 in the lunar exosphere. Geophys Res Lett 43:22–27 Hoffman JH, Hodges Jr RR, Johnson FS, Evans DE (1973) Lunar atmospheric composition results from Apollo 17. In: Lunar and planetary science conference proceedings, Houston, vol 4, p 2865 Hurley DM, Cook JC, Benna M, Halekas JS, Feldman PD, Retherford KD, Hodges RR, Grava C, Mahaffy P, Gladstone GR, Greathouse T, Kaufmann DE, Elphic RC, Stern SA (2016) Understanding temporal and spatial variability of the lunar helium atmosphere using simultaneous observations from LRO, LADEE, and ARTEMIS. Icarus 273:45–52 Kegerreis JA, Eke VR, Massey RJ, Beaumont SK, Elphic RC, Teodoro LF (2017) Evidence for a localized source of the argon in the lunar exosphere. J Geophys Res 122:2163–2181 Killen RM (2002) Source and maintenance of the argon atmospheres of Mercury and the Moon. Meteorit Planet Sci 37:1223–1231 Lawson SL, Feldman WC, Lawrence DJ, Moore KR, Elphic RC, Belian RD, Maurice S (2005) Recent outgassing from the lunar surface: the Lunar Prospector Alpha Particle Spectrometer. J Geophys Res 110: E09009 Stern SA (1999) The lunar atmosphere: history, status, current problems, and context. Rev Geophys 37:453–492 Wieler R, Heber VS (2003) Noble gas isotopes on the Moon. Space Sci Rev 106:197–210

Non-mare Volcanism ▶ Silicic Volcanism on the Moon

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Opposition Effect Brian Cudnik Department of Chemistry and Physics, Prairie View A&M University, Prairie View, TX, USA

Definition The opposition effect (also known as opposition surge) is the brightening of a rough surface or an object made up of many particles when illuminated by a source located directly behind the observer.

Theory and Application: Physical Causes and Mechanisms The opposition effect (Fig. 1) is often referred to by other names, e.g., the opposition surge, opposition spike, or Seeliger effect. It describes the phenomenon where an object or region noticeably brightens when that object or region reaches a phase angle of approximately zero. This effect has been known for over a century, being first observed in the reflected sunlight from Saturn’s rings at opposition (Seeliger 1895; Hapke et al. 1998). This phenomenon was first observed with the Moon by Gehrels et al. (1964) as the brightness of the lunar surface is almost 40% greater at the moment of Full Moon than 1 day before or after Full Moon (Hapke et al. © Springer Nature Switzerland AG 2023 B. Cudnik (ed.), Encyclopedia of Lunar Science, https://doi.org/10.1007/978-3-319-14541-9

1998). The first laboratory experiments to observe the opposition effect in powders were performed by Oetking (1966). Early studies of the lunar opposition effect were done using Apollo photographs by Whitaker (1969) and Pohn et al. (1969) (Hapke et al. 2012), but it was not until Clementine that this phenomena was studied with lunarorbiting spacecraft. The phase angle is defined as the angle from the Sun, to an object, to an observer, with the object at the vertex of the angle. The phase angle is essentially the quantification of the apparent phase of a moon or planet and is defined such that an 180
phase angle corresponds to the New Moon, a phase angle of 90
indicates a quarter Moon, and a 0
phase angle corresponds to a Full Moon (Lakdawalla 2009). In the last case, the Sun is directly behind the observer and the object is directly in front of the observer. The opposition effect is typically observed at phase angles less than approximately 10
and was thought for most of the time since its initial discovery to result solely from the hiding of the shadows of the rocks and other irregularities in the terrain. This phenomenon is known as shadow hiding, and this occurs when all shadows, from the observer’s perspective, disappear, which leads to a sudden and often dramatic increase in brightness. This increase is the result of undiminished light being reflected back to the observer. The Moon, as an airless object, demonstrates this effectively due to a lack of atmosphere (which scatters and diffuses light),

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Opposition Effect, Fig. 1 The opposition surge that encircled the head of Apollo 11 astronaut Buzz Aldrin. This represents the typical appearance of an opposition surge. (Image courtesy of NASA)

as well as regolith abundance. On a regolith surface, sunlight illuminates the small irregularities, bumps, pits, and other defects, which normally lie in the shadow, thereby adding to the brightness of the scene (Seeliger 1887). Compared to other airless planetary bodies, the Moon displays a broad opposition curve profile (Shkuratov et al. 1999). It has been widely thought that shadow hiding is the dominant mechanism behind the lunar opposition surge phenomena, owing mainly to the Moon’s low albedo. However, work by Hapke et al. (1998) and others have demonstrated that this may not be the case. With spacecraft-based observations and ground-based experiments involving lunar regolith samples, it became clear that another process was involved in producing the observed opposition effect. For phase angles less than 1–2
, an extremely narrow and large surge takes place that results from the constructive interference of backscattered light with incoming light. In theory, this adds to the opposition effect observed with the Moon. This additional intensity surge is known as coherent backscatter and is seen not only with the Moon but also in a number of solar

Opposition Effect

system objects. For coherent backscatter to work best, the size of scatter in the Moon’s regolith is comparable to the wavelength of light and the distance between scattering particles is greater than one wavelength of visible light (e.g., 550 nm, Hapke 1989). It is presently known that one effect or the other alone does not account for the observed profile of the opposition effect; both processes are involved. Opposition effect is useful and important in planetary science such that a careful analysis of the curve of the effect yields data on the state of compactness of the surface as well as size distribution of the particles on the surface producing the surge (Hapke 1986; Buratti et al. 1996). Regolith properties of objects other than the Moon are interpreted in terms of our models of the lunar regolith, since this is the only other world, besides Earth, that we have samples of regolith (Hapke et al. 1998), with the exception of small bits of the Itokawa asteroid returned by Hayabusa in 2010. The angular width of the curve produced by the shadow-hiding opposition effect is independent of wavelength, but a proportional dependence of wavelength is noted for the coherent background opposition effect (Hapke et al. 1998), as described by an expression for half width at half maximum (HWHM) of the coherent backscatter effect given by (Van Albada et al. 1990), HWHM ¼ bl=2pL,

ð1Þ

with l representing the wavelength of light, b is a constant equal to 0.36, and L is the transport mean free path for the photons in the medium. L is given by L ¼ ½nσQs ð1–½ cos θÞ−1

¼ ½nσwQE ð1–½ cos θÞ−1

ð2Þ

With n being the number of particles per unit volume, s signifying the mean particle geometrical cross section, Qs is the mean scattering efficiency, and QE is the mean extinction efficiency. The mean single scattering albedo is

Opposition Effect

937

given by the ratio w ¼ Qs/QE, and the average cosine of the scattering angle is depicted by [cos θ], and may vary with wavelength.

Observations and Results

Opposition Effect, Fig. 2 The integral phase curve of the Moon at three wavelengths. (Plot adapted from Fig. 5 of Buratti et al. 1996, used with permission)

Normalized, disk–integrated brightness

Various research teams have investigated the opposition surge phenomena over the years. Bonnie Buratti and her team used multispectral observations of the opposition surge, recorded by the Clementine satellite in 1994. They have found that there is an increase in brightness of approximately 40% between a phase angle of 0
and 4
(Fig. 2). This increase varies with surface roughness, i.e., it is greater for the rugged Highlands compared with the much smoother Maria. Results from Buratti et al. (1996) show that the opposition surge has only a minor wavelength dependence, i.e., 3–4% larger at 0.41 mm than at 1.00 mm. This suggests that the dominant mechanism associated with surge arises from shadow hiding rather than coherent backscatter (Buratti et al. 1996). Only 10% of the opposition surge amplitude significantly depends on terrain: the Highlands experience a 10% increase in surge due to textural variations between the rougher and brighter Highlands compared with the flatter, smoother, and darker Maria. Observations of the Moon at phase angles less than one degree are not possible from the Earth’s surface due to the Earth’s own shadow. One may see the intensity of the Full Moon

increase as it approaches zero phase, but it starts to become eclipsed by Earth’s umbra as it moves to within a degree or so of zero phase. In order to observe the Moon at zero phase, it is necessary to do so by spacecraft. In fact, the Clementine spacecraft obtained several hundred images of the opposition effect from 0.41 to 2.8 microns. This mission provided the first multispectral observations of the lunar opposition effect at phase angles less than a few degrees. The results enabled researchers to obtain highly accurate phase curves within a few degrees of zero phase angle. Further analyses, from Buratti et al. (1996) with the Clementine data along with telescopic lunar data, showed the shadow-hiding mechanism to be the primary cause of the lunar opposition surge. They acknowledge that coherent backscatter is not entirely absent; work by Hapke et al. (1993) with laboratory measurements of Apollo samples down to 1
phase angle indicate the presence of coherent backscatter as revealed by the polarization signature. Furthermore, Hapke et al. (1998) asserts that their data can be interpreted in other ways. For instance, photometric differences between the Highlands and Maria support coherent backscattering, as well as shadow hiding. Polarization of the incident light also has an effect on the widths of the peak of the coherent backscatter opposition effect (Van Albada et al. 1987). Hapke et al. (1998) performed laboratory studies on eight samples of mature and immature regolith from the Highlands and Maria, using

1.1 1.0 0.9 0.8 0.7 0.6 0.5 0.4 0.41 mm 0.75 mm 1.0 mm

0.3 0.2 0.1 0.0 0

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40

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80

100

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Solar phase angle

140

160

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Opposition Effect

depicted in the plot indicates again that shadow hiding is the dominant factor. Lunar soil is reddish which results in albedo increasing with increasing wavelength. If the correlation were positive, then coherent backscattering would be the dominant mechanism at play but this is not what is observed. Figure 5 is a plot from the same group showing the ratio of total reflectances of the lunar samples measured in red light to that in blue light. If the opposition effect were entirely due to shadow hiding, then the opposition peak width should be independent of wavelength and albedo. The red/blue ratio would decrease as the phase angle decreases. If coherent backscattering were dominant, then the peak in amplitude and width

blue laser light at 442 nm and red laser light at 633 nm. They studied the effects of polarization on two of the eight lunar samples they worked with and measured both circularly and linearly polarized light (blue and red) reflected from the samples. If the opposition effect were primarily due to coherent backscattering, then the angular width of reflectance should be notably wider in the red and in the blue, but Fig. 3 shows there is little difference, indicating that shadow hiding does indeed dominate in the case of lunar sample 10,084. Figure 4 depicts the amplitude of the opposition effect of all eight samples studied by Hapke et al. (1998) and their negative correlation Opposition Effect, Fig. 3 Total reflectances, normalized at 10, of the lunar soil sample 10084 measured in red and blue light versus phase angle. (This is Fig. 6 from Hapke et al. (1998), used with permission)

1

Normalized Reflectance

Lunar Soil 10084 0.9

0.8

0.7

Red

0.6

0.5

0

2

6 4 Phase Angle (degrees)

8

10

2.6 Blue Red

2.4 2.2 I (1°) /I (10°)

Opposition Effect, Fig. 4 Amplitude of the opposition effect of all eight of the lunar samples investigated by Hapke et al. (1998). The samples were measured in red and blue light and these measurements are plotted versus normal albedo. The lines connect the pair of points for the same sample, eight lines in all. (This is Fig. 7 from Hapke et al. (1998), used with permission)

Blue

2 1.8 1.6 1.4 1.2 1 0

5

10

15 20 25 Normal Albedo (%)

30

35

40

Opposition Effect

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Opposition Effect, Fig. 5 Ratios of the reflectances in red light to those of blue light versus phase angle for all eight samples studied; ratios are normalized at 10o. (This is Fig. 8 from Hapke et al. (1998), used with permission)

would both increase with increasing albedo. This would show in the plot as an increase of the red/blue ratio as the phase angle decreases. The plot shows that the ratio decreases down to phase angles of 3
to 4
, suggestion shadow hiding dominates for phase angles greater than 3
; the ratio then increases at smaller phase angles, indicating that coherent backscattering becomes an important contributor to the opposition surge. Neither coherent backscattering nor shadow hiding alone account for the opposition surge, but it is possible that each have roughly equal contributions. Both mechanisms are needed to reproduce what is actually observed on the Moon. The Full Moon in opposition effect does not show neither limb darkening nor limb brightening which further supports this conclusion. Further, Hapke et al. (1998) found a narrow coherent backscattering peak (half-width of about 20) that was superimposed on a broader, shadow-hiding peak (about 80 wide); both peaks had similar amplitude. The consensus, in this case, is that both causes, i.e., shadow hiding and coherent backscatter, are roughly equally responsible for lunar opposition surge. In the case of a Full Moon, we observe a nearly 40% brightness surge versus the day before and after the Full Moon. Coherent scattering is important for remote sensing because the opposition surge profile provides information on surface porosity through the transport mean free path of photons in regoliths.

The Wide Angle Camera of the Lunar Reconnaissance Orbiter made observations in seven narrowband wavelengths between 321 and 689 nm, and at phase angles ranging from 0
to 120
. Hapke et al. (2012) analyzed the Highlands area and found that coherent backscatter contributed nearly 40% of the opposition surge in the UV, increasing to over 60% in the red. Their conclusion is supported by laboratory work they did earlier on samples of lunar regolith by measuring their circular polarization ratios. They also show that “phase reddening is caused by the increased contribution of interparticle multiple scattering as the wavelength and albedo increase.”

Modeling the Opposition Surge In addition to the observations and laboratory analyses highlighted above, several studies have undertaken the mathematical modeling of the opposition surge phase function (Hapke et al. 2012). Hapke’s phase function is summarized by the following equation (Velikodsky et al. 2016): pðaÞ ¼ ðPðaÞ  ð1 þ f s ðaÞÞ þ Mði, eÞÞ  ð 1 þ f c ð aÞ Þ

ð3Þ

where P(α) is the average single particle scattering function, M (i, e) is the contribution from interparticle multiple scatterers, and fs (α)

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and fc (α) are the shadow and coherent backscatter functions in the opposition surge, respectively. This model is unable to separate the two functions. The current photometric model used by Akimov (Velikodsky et al. 2016) allows one to separate fs (α) and fc (α) but is not the only such model available. The opposition spike that arises from the coherent backscattering opposition effect is contained in the following phase function: pðaÞ ¼ expðmaÞ  f ðaÞ

ð4Þ

where m is the phase curve slope over a wide range of phase angles and f (α) is the additional brightness surge component near opposition. The exp. (mα) term describes the shadowing effect as the phase angle approaches zero. Hapke et al. (2012) performed a quantitative analysis on Lunar Reconnaissance Orbiter (LRO) data for the lunar phase curve out to phase angles of 120
. This included the opposition surge, which they modeled in their analysis. Their entry includes the full treatment of this analysis, which takes into account both the shadow hiding and coherent backscattering opposition effects; the “highlights” are summarized in this entry. The shadow-hiding opposition effect is quantified by the angular shape function Bs ðgÞ ¼ 1=½1 þ tanðg=2Þ=hs 

ð5Þ

with g representing the phase angle and hs the shadow-hiding opposition effect angular width parameter. The coherent backscatter opposition surge is quantified as follows BC ð g Þ ¼

1 þ ½1  expðzÞ=z , 2½ 1 þ z  2

¼ tanðg=2Þ=hc

with z ð6Þ

where hc being the coherent backscatter opposition effect angular width parameter. Hapke et al. (2012) continues in the explanation of the source of Eq. (6) and how this was modified slightly in his earlier paper (Hapke 2002). The angular width parameter of each opposition effect is then

correlated to the properties of the lunar regolith as described by Hapke et al. (1993) for Eq. (7) and Akkermans and Maynard (1986) for Eq. (8) hs ¼ hai=2LE

ð7Þ

hc ¼ l=4pLT

ð8Þ

where hai is the mean particle radius, ΛE is the extinction mean free path, l is the wavelength, and ΛT is the transport mean free path (the distance a photon travels before being scattered by a large angle). These equations are widely accepted but may not be valid in certain situations as described in section 5 of Hapke et al. (2012). The half-width at half maximum of the two opposition effects are related to hs and hc by Eqs. (9) and (10) gs ¼ 2hs ¼ hai=LE

ð9Þ

gc ¼ 0:72hc ¼ 0:36l=2pLT

ð10Þ

and the amplitude of the shadow-hiding effect is Bs ¼ Rð0Þ=wpð0Þ

ð11Þ

where R(0) is the amount of light backscattered “into zero phase angle from the fraction of the particle cross-sectional area at or near the surface of a typical particle that is directly illuminated by the incident light.” Hapke et al. (2012). Again, the reader is encouraged to consult this and other references therein for more details on the theoretical discussion of opposition effect and phase curve photometry. The capabilities of the LRO Wide Angle Camera enabled both types of opposition effects to be separated out in the data. Hapke et al. (2012) conclude that the coherent backscatter accounts for nearly 40% of the opposition effect in the UV part of the spectrum. This percentage increases to over 60% in the red part of the spectrum. Although this conclusion has been derived from only one Highlands location on the Moon, according to this research group, it is likely that this conclusion applies to most places on the Moon.

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Buratti et al. (1996) modeled the opposition curves based on their shadow-hiding model as functions of single-scattering albedo (w) and plotted the curve for values of w ranging from 0.1 to 0.9 (Figs. 6 and 7). They then fit the standard model to actual Clementine satellite data (Fig. 4) using five different values for the opposition surge width parameter, h. Velikodsky et al. (2016) used LROC WAC data and studied the opposition surge for both the Maria and Highlands. They found a coherent backscattering component in the opposition surge within 1.2
of the opposition for the Highlands in red light and within 3.9
for the Maria in blue light. The general slope of the lunar phase curve is indicative of the average particle phase function, shadow hiding, and incoherent multiple scatterers. Overall, the opposition surge has an amplitude of up to 8% and a width of a few degrees. Increasing albedo appears to lead to an increase in amplitude with a decrease in width (Velikodsky et al. 2016). The amplitude for the Moon is much smaller than that for bright surfaces, such as

those on icy satellites located in the outer solar system. A noticeable component of the coherent backscatter is an enhancement at α < 0.25
. This is possibly due to interparticle scattering in the lunar regolith. In summary, the opposition effect varies between the different types of lunar surfaces due to variations in albedo. Immature lunar soils have very narrow opposition surges due to larger particles that are more homogeneous. Despite the wealth of information from several research groups summarized above, the coherent backscatter phenomenon has not been extensively studied. From ground-based observatories, there is no reliable method available to study contributions from coherent backscatter primarily due to the fact that the Earth’s shadow obscures the Moon when the phase angle is less than 2
. Earth-based spectrophotometric observations are not yet accurate enough to reliably detect variations in the opposition effect over the entire lunar surface, and only a handful of spacecraft so far have produced the data needed to enable such studies.

1 w=0.1

w=0.5

w=0.2

w=0.7

w=0.3

w=0.9

Normalized Reflectance

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1

2

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Phase Angle (°) Opposition Effect, Fig. 6 The predicted opposition curves based on the shadow-hiding model as a function of single scattering albedo (w). Note that the effect that

albedo has on the phase curves becomes significant when w is greater than 0.5. (Plot is Fig. 6 of Buratti et al. 1996, used with permission)

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1.05

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1 0.95

Clementine Data

h = 0.09

Standard Model, h= 0.044

h = 0.12

Standard Model, h= 0.12

h = 0.15

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0.5

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2.5

3

3.5

4

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5

Phase Angle (°) Opposition Effect, Fig. 7 The fits of standard (dotted and solid lines) and two-layer (remaining lines) shadowhiding models to Clementine data at small solar phase angles. The fit is shown for four different values of Hapke’s (1986) opposition surge width parameter, h. For both layers, w ¼ 0.245 and the particle size ratios in the two

layers are 0.64, 0.29, 0.19, and 0.14; these correspond with values of h that range from lowest to highest. The model takes into account the finite size of the Sun as seen from the Moon’s surface. (Plot is Fig. 7 of Buratti et al. 1996, used with permission)

Summary

primarily of smooth volcanic plains. The former produces an approximately 10% larger surge than the latter. Hapke et al. (2012) also found that, with LROC data, the coherent backscatter component of the opposition effect ranges from nearly 40% in the UV to over 60% in the red region of the spectrum. This team also showed that the angular width of the opposition effect is nearly independent of wavelength, contrary to theories that predicted a trend that follows the square of the wavelength. Part of this wavelength independence may be the result of phase reddening “caused by the increased contribution of interparticle multiple scattering as the wavelength and albedo increase” (Hapke et al. 2012). Various groups have quantitatively modeled the opposition effect and its various features and components. Velikodsky et al. (2016) found that the coherent backscattering opposition effect displays a small amplitude of up to 8% and a width of a few degrees, in visible light. Increase in lunar surface albedo leads to an increase in amplitude and a decrease in width, in agreement

The opposition effect, or opposition surge as it is also called, refers to the dramatic brightening of a surface at and near the subsolar point. This is commonly observed throughout the solar system in a variety of settings. The lunar regolith provides a prototype to study the electromagnetic scattering properties of regoliths from other worlds. Two physical mechanisms are thought to be at play during this effect: shadow-hiding and coherent backscattering. Shadow hiding occurs when shadows cast by particles on the lunar surface disappear due to the process where each particle hides its own shadow. Coherent backscattering happens when two parts of a wave travel in the same, multiply scattered path through the regolith but in opposite directions, which constructively interferes with phase angles near zero (Hapke et al. 2012). For the Moon, the opposition effect varies between the Highlands, which consist of rougher terrain, and the Maria, which is composed

Optical Maturity (OMAT)

with models. Different types of lunar surface can exhibit different widths of the opposition effect (specifically the coherent backscatter component) due to variations in albedo. Immature lunar soils display a very narrow (10 million-year timescales due to steady decline of the Moon’s axial tilt with increasing Earth-Moon separation (Ward 1975). The prevalence of PSRs increases poleward as topography is more easily able to cast larger shadows due to higher solar incidence angles at high latitudes. Many polar craters have permanently shadowed interiors due to their concave topography, with the crater rim shielding the wall and floor from direct illumination by sunlight on all sides. The predominant sources of radiation in PSRs are limited to upwelling heat flow from the lunar interior, stellar radiation, secondary illumination by reflected or scattered light from illuminated surfaces, or thermal emission from nearby warmer shadowed surfaces. These radiative sources are often negligible, allowing PSRs to remain at very low temperatures that are more akin to those observed on outer solar system

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Surface and Near-Surface Thermal Environment of the Moon

bodies. Major PSRs are clearly delineated from surrounding terrain that is not permanently shaded by mapping the maximum observed temperature. Figure 3 illustrates the maximum bolometric temperature (see section “Bolometric Brightness Temperature”) observed at the lunar south pole using LRO Diviner data. Lunar cold traps were first proposed by Watson et al. (1961) to be areas in permanent shadow where surface temperatures do not exceed 100 K, causing any volatiles to become trapped by lack of available energy for mobility. The source for such volatiles is unclear, but could be ejecta from impacting comets or compounds formed by the implanting of solar wind H+ ions. Removal of surface volatiles may occur via solar wind sputtering, sublimation, or ejection by impacts. Once deposited, the permanence of coldtrapped surface volatiles depends on their shielding from loss processes. While surface water ice has been shown to be stable over million-year timescales in many locations at the lunar south pole (Paige et al. 2010b), diffusion to depth of volatiles would provide protection over geologic time, allowing long-term stability of subsurface deposits. 150 6

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240

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220 2

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5

0

3

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Distance from South pole (km)

Surface and NearSurface Thermal Environment of the Moon, Fig. 3 Maximum bolometric temperature at the lunar south pole as observed by the Diviner Lunar Radiometer. Major permanently shadowed regions (PSRs) and the impact site of the LCROSS mission are marked (Data used in this figure were acquired between July 2009 and September 2013)

Evidence for surface water ice deposits in PSRs has been presented as anomalies in albedo (Riner et al. 2013) and suppression of neutrons inferred to be by ice (Mitrofanov et al. 2010); however, the case for surface water ice deposits in lunar PSRs remains debated, particularly when compared to similar but much stronger evidence for polar water ice deposits discovered on Mercury (Paige et al. 2013). The LCROSS mission, launched jointly with the LRO mission in 2009, involved the impact of a spent Centaur rocket stage into a south polar permanently shadowed region near Cabeus crater (location 7 on Fig. 3). The impact generated a plume of ejecta which was observed by nine instruments aboard a shepherding spacecraft, as well as instruments aboard NASA’s Lunar Reconnaissance Orbiter. Absorbance in the near infrared was attributed to water vapor and ice, while ultraviolet emissions were thought to be due to hydroxyl radicals, which further supported the presence of water in the debris cloud. A total regolith abundance of water ice in the regolith at the impact site was determined to be 5.6  2.9 % by mass. While water was the most abundant volatile detected by a factor of roughly 6, spectral

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Haworth Shoemaker Faustini Shackleton deGerlache Cabeus LCROSS impact

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200 180 160 140 120 100 80 60

Surface and Near-Surface Thermal Environment of the Moon

fits indicated that a wide inventory of subsurface volatiles were present in the plume. Moderate to minor amounts of several other volatile species were detected, including (in decreasing order of abundance) H2S, NH3, SO2, C2H4, CO2, CH3OH, CH4 and OH (Colaprete et al. 2010). While a full inventory of cold-trapped subsurface volatiles in PSRs remains unquantified, direct observation of volatiles in the LCROSS ejecta plume suggests that a wide range of volatiles are stable in subsurface deposits in PSRs. Heat Flow The subsurface temperature on the Moon, as for many airless bodies in the solar system, is governed by the insolation-radiation balance at the surface and upwelling heat flow from the interior, produced by radioactive decay of U, Th, and K. At depths beyond the influence of diurnal and seasonal thermal waves, the vertical temperature gradient dT/dz is controlled by vertical heat flux from the lunar interior Fz. These quantities are related to the thermal conductivity, k, by the relationship Fz ¼ k

dT dz

ð6Þ

The magnitude of radiogenic heat flow at the surface and its spatial variation can be used to determine the interior composition, structure, and evolution of the Moon. The Apollo Heat Flow Experiment on board the Apollo 15 and 17 missions obtained two measurements of 21  3 mW m2 and 15  2 mW m2 (Langseth et al. 1972) at their respective landing sites. In the absence of any in situ measurements since, the large difference between these two values has prompted debate over their meaning with regard to the radiogenic properties of the lunar interior. A recent reanalysis and modeling study of heat flow at the Apollo landing sites (Siegler and Smrekar 2014) found that surface heat flow can be greatly influenced by deep subsurface radiogenic content and crustal density but that overall crustal thickness and the presence of radiogenicrich impact ejecta have a comparatively small effect.

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Cold Spots Thermal “cold spots” are documented as regions around fresh impact craters that are unusually cold at night, i.e., that cool faster than their surroundings when not illuminated and therefore likely have low thermal inertias. Regions extend radially up to hundreds of crater radii from the central craters and over 400 cases of this phenomenon have been documented. This material is unlikely to be crater ejecta due to the large volumes of material that would be required. A current working hypothesis is that these regions have been subject to in situ decompression of regolith, perhaps due to turbulent vapor or scouring by ballistic particles (Bandfield et al. 2014b). This is in contrast to higher thermal inertias observed over much smaller distances from impact craters, due to the increased abundance of blocky ejecta material. Surface Rock Abundance The thermal effect of massive blocky material on the lunar surface is to increase the apparent thermal inertia of the region, because cohesive dense boulders have higher thermal inertias than average lunar regolith. Observations of rocky areas therefore show lower daytime and higher nighttime brightness temperatures. Rock abundance is higher in impact crater ejecta blankets, but abates with time as larger rocks are broken down by micrometeorite impacts. Rock abundance and impact crater age are therefore well correlated (Bandfield et al. 2011), and this rate of rock breakdown or “regolith production rate” has only recently begun to be constrained (Ghent et al. 2014; Hayne et al. 2013). Rocks on the lunar surface are best preserved on young surfaces or on steep slopes, where mass wasting prevents mantling with fines.

References Bandfield JL, Ghent RR, Vasavada AR, Paige DA, Lawrence SJ, Robinson MS (2011) Lunar surface rock abundance and regolith fines temperatures derived from LRO Diviner Radiometer Data. J Geophys Res 116:E00H02

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Bandfield JL, Hayne PO, Paige DA (2014a) What is the surface temperature of the moon? In: 45th lunar and planetary science conference no 1519. Lunar and Planetary Institute, The Woodlands, Houston, TX, USA Bandfield JL, Song E, Hayne PO, Brand BD, Ghent RR, Vasavada AR, Paige DA (2014b) Lunar cold spots: granular flow features and extensive insulating materials surrounding young craters. Icarus 231:221–231 Colaprete A, Schultz P, Heldmann J, Wooden D, Shirley M, Ennico K, Hermalyn B, Marshall W, Ricco A, Elphic RC, Goldstein D, Summy D, Bart GD, Asphaug E, Korycansky D, Landis D, Sollitt L (2010) Detection of water in the lcross ejecta plume. Science 330:463–468 Gault DE, Horz F, Brownlee DE, Hartung JB (1974) Mixing of the lunar regolith. In: Proceedings of the fifth lunar science conference, vol 3, Lunar Science Institute, Houston, TX, USA, pp 2365–2386 Ghent RR, Hayne PO, Bandfield J, Campbell BA, Allen CC, Carter LM, Paige DA (2014) Constraints on the recent rate of lunar ejecta breakdown and implications for crater ages. Geology 42(12):1059–1062 Greenhagen BT, Lucey PG, Wyatt MB, Glotch TD, Allen CC, Arnold JR, Bandfield JL, Bowles NE, DonaldsonHanna KL, Hayne PO, Song E, Thomas IR, Paige DA (2010) Global silicate mineralogy of the moon from the diviner lunar radiometer. Science 329:1507–1509 Hapke B, van Horn H (1963) Photometric studies of complex surfaces, with applications to the moon. J Geophys Res 68(15):4545–4570 Hayne PO, Greenhagen BT, Siegler MA, Vasavada AR, Aharonson O, Bandfield JL, Ghent RR, Elphic RC, Paige DA (2011) The moon’s extremely insulating near-surface: diviner infrared observations of a total lunar eclipse. In: Fall meeting no P13D-1712. American Geophysical Union, Moscone Center, San Francisco, CA, USA Hayne PO, Ghent R, Bandfield JL, Vasavada AR, Siegler MA, Greenhagen BT, Williams J-P, Paige DA (2013) Formation and evolution of the moon’s upper regolith: constraints from diviner thermal measurements. In: 44th lunar and planetary science conference no 3003. Lunar and Planetary Institute, The Woodlands, Houston, TX, USA Helfenstein P, Shepard MK (1999) Submillimeter-scale topography of the lunar regolith. Icarus 141:107–131 Langseth MG, Clark SP, Chute JL, Keihm J, Wechsler AE (1972) The apollo 15 lunar heat-flow measurement. In: Lamont-Doherty geological observatory contribution number 1800, Conference on Lunar Geophysics, Lunar Science Institute, Houston, TX, USA Lawson SL, Jakosky BM, Park H-S, Mellon MT (2000) Brightness temperatures of the lunar surface: calibration and global analysis of the Clementine long-wave infrared camera data. J Geophys Res 105(E2):4273–4290 Mitrofanov IG, Sanin AB, Boynton WV, Chin G, Garvin JB, Golovin D, Evans LG, Harshman K, Kozyrev AS, Litvak ML, Malakhov A, Mazarico E, McClanahan T,

Milikh G, Mokrousov M, Nandikotkur G, Neumann GA, Nuzhdin I, Sagdeev R, Shevchenko V, Shvetsov V, Smith DE, Starr R, Tretyakov VI, Trombka J, Usikov D, Varenikov A, Vostrukhin A, Zuber MT (2010) Hydrogen mapping of the lunar south pole using the lro neutron detector experiment lend. Science 330:483–486 Paige DA, Foote MC, Greenhagen BT, Schofield JT, Calcutt SB, Vasavada AR, Preston DJ, Taylor FW, Allen CC, Snook KJ, Jakosky BM, Murray BC, Soderblom LA, Jau B, Loring S, Bulharowski J, Bowles NE, Thomas IR, Sullivan MT, Avis C, De Jong EM, Hartford W, McCleese DJ (2010a) The lunar reconnaissance orbiter diviner lunar radiometer experiment. Space Sci Rev. doi:10.1007/s11214-0099529-2 Paige DA, Siegler MA, Zhang JA, Hayne PO, Foote EJ, Bennett K, Vasavada AR, Greenhagen BT, Schofield JT, McCleese DJ, Foote MC, DeJong E, Bills BG, Hartford W, Murray BC, Allen CC, Snook K, Soderblom LA, Calcutt SB, Taylor FW, Bowles NE, Bandfield JL, Elphic RC, Ghent R, Glotch TD, Wyatt MB, Lucey PG (2010b) Diviner lunar radiometer observations of cold traps in the moon’s south polar region. Science 330:479–482 Paige DA, Siegler MA, Harmon JK, Neumann GA, Mazarico EM, Smith DE, Zuber MT, Harju E, Delitsky ML, Solomon SC (2013) Thermal stability of volatiles in the north polar region of mercury. Science 339(6117):300–303 Putzig NE, Mellon MT, Kretke KA, Arvidson RE (2005) Global thermal inertia and surface properties of mars from the mgs mapping mission. Icarus 173:325–341 Riner MA, Lucey PG, Neumann GA, Smith DE, Zuber MT, Bussey DBJ, Cahill JTS, Mazarico EM (2013) Albedo of permanently shadowed regions (psrs) at the lunar south pole. In: 44th lunar and Planetary science conference no 2677, The Woodlands, Houston, TX, USA Sefton-Nash E, Williams J-P, Paige DA (2014) Modeling, gridding and storage of effective fields-of-view for terascale, point-based planetary datasets: case study – lro diviner. In: 45th lunar and planetary science conference no 2737. Lunar and Planetary Institute, The Woodlands, Houston, TX, USA Siegler MA, Smrekar SE (2014) Lunar heat flow: regional prospective of the Apollo landing sites. J Geophys Res Planets 119:47–63 Spencer JR (1990) A rough-surface thermophysical model for airless planets. Icarus 83:27–38 Vasavada AR, Paige DA, Wood SE (1999) Near-surface temperatures on mercury and the moon and the stability of polar ice deposits. Icarus 141:179–193 Ward WR (1975) Past orientation of the lunar spin axis. Science 189(4200):377–379 Watson K, Murray BC, Brown H (1961) The behavior of volatiles on the lunar surface. J Geophys Res 66(9):3033–3045

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Lunar Rocks

Surface of the Moon, Distribution of Materials and Structures Shengbo Chen College of Geo-Exploration Science and Technology, Jilin University, Jilin, China

Introduction The surface of the moon is a critical window to understanding the entire moon using remote sensing. The moon surface consists of a layer of lunar regolith covering the underlying bedrock. Regolith is the layer or mantle of fragmented and unconsolidated rock material potentially residual or in transit and highly variable that forms the surface nearly everywhere (Jolliff et al. 2006). In addition to the interaction of high energy particles, the major physical changes that occur with lunar regolith maturation include comminution and agglutination between rocks and minerals. At the same time, the lunar structure that is formed by tectonic activities can lead to the uneven distribution of chemical elements and can provide information about rocks and minerals in the subsurface of lunar crust. In the end, the distribution and composition of lunar surface material can be described by four processes: petrology, mineralogy, element chemistry, and structural geology. In addition, water has attracted much attention from the lunar science community for 40 years in the hope of exploring this resource and learning more about the moon’s origin. Our knowledge of the lunar interior stems from the information of minerals and elements on the lunar surface coupled with the lunar geological structure. Thus, understanding the composition and distribution of lunar surface materials (rocks, minerals, elements, and ice water) helps reconstruct the internal moon structure and has vital significance to recognizing its origin and evolution.

Traditionally, the lunar terrain has been divided into the highlands and the maria based on the smoothness and brightness of the lunar surface. The corresponding dominant rock types are the highland anorthosite and mare basalt, respectively (Ling et al. 2014). Mare Basalt Mare basalts are distributed in the lunar maria and are characterized by the large, dark, basaltic plains on the lunar surface. Maria are much darker than the highlands, occupy more than 17% of the lunar surface, and are concentrated on the lunar nearside. Mare basalts are formed from ancient volcanic eruptions and basaltic magma filling large impact basins. The thickness of mare basalts is normally from 500 m to 1,300 m and up to 4,500 m in a few large basins. Mare basalts have been detected in many lunar samples returned from the Apollo and Luna missions. The major minerals in mare basalt are pyroxene and plagioclase but also include olivine and ilmenite in varying concentrations. Mare basalts are rich in FeO and TiO2, but poor in Al2O3 and CaO. The content of TiO2 varies widely from 0.3% to 13%. Accordingly, mare basalts are usually divided into three major classes based on the different TiO2 contents that are highTi mare basalts (>9 wt.% TiO2), low-Ti mare basalts (1.5–9 wt.% TiO2), and very-low-Ti (VLT) mare basalts (1 m across),

surrounding the base of the Schrödinger central peak ring. Yellow circles indicate large boulders showing macrofracturing. Image width 416 m, north is up. (Image credit: NASA/GSFC/Arizona State University)

Damage Effects The morphology of boulders on the Moon can exhibit a range of fractured textures (Ruesch et al. 2020). Documenting these textures based on fragment size, and the original boulder size, and orientation of the fractures can give us insight into the local boulder fragmentation processes. Figure 1 shows an example of a boulder field near the central peak ring of Schrödinger crater. From these textures, quantitative analyses on the fragment spatial density can be observed (Ruesch et al. 2020). Higher spatial density of fragmented

Thermophysical Behavior of the Lunar Surface

boulders typically consists of a large fragment and smaller fragments in the perimeter of the parent boulder. Lower spatial density would have a relatively large number of smaller fragment pieces and sometimes more dispersed (Ruesch et al. 2020).

Cross-References ▶ Lunar Crater Ejecta ▶ Lunar Rocks ▶ Lunar Surface, Interaction of the Solar Wind with Upper Regolith ▶ Regolith Structure

References Attewell PB, Farmer IW (1973) Fatigue behaviour of rock. In: International journal of rock mechanics and mining sciences & geomechanics abstracts, vol 10, No. 1, Pergamon, pp 1–9 Basilevsky AT, Head JW, Horz F (2013) Survival times of meter-sized boulders on the surface of the Moon. Planet Space Sci 89:118–126 Dong Y, Tagavi KA, Lovell MR, Deng Z (2000) Analysis of stress in cross wedge rolling with application to failure. Int J Mech Sci 42(7):1233–1253 Eppes MC, Magi B, Hallet B, Delmelle E, MackenzieHelnwein P, Warren K, Swami S (2016) Deciphering the role of solar-induced thermal stresses in rock weathering. Bulletin 128(9–10):1315–1338 Ghent RR, Hayne PO, Bandfield JL, Campbell BA, Allen CC, Carter LM, Paige DA (2014) Constraints on the recent rate of lunar ejecta breakdown and implications for crater ages. Geology 42(12):1059–1062 Kranz RL (1983) Microcracks in rocks: a review. Tectonophysics 100(1–3):449–480 Levi FA (1973) Experimental evidence of low-temperature thermal metamorphism in stony meteorites. Meteoritics 8:408–409 Molaro J, Byrne S (2012) Rates of temperature change of airless landscapes and implications for thermal stress weathering. J Geophys Res Planets 117(E10) Molaro JL, Byrne S, Langer SA (2015) Grain-scale thermoelastic stresses and spatiotemporal temperature gradients on airless bodies, implications for rock breakdown. J Geophys Res Planets 120(2):255–277 Molaro JL, Byrne S, Le JL (2017) Thermally induced stresses in boulders on airless body surfaces, and implications for rock breakdown. Icarus 294:247–261 Ravaji B, Alí-Lagoa V, Delbo M, Wilkerson JW (2019) Unraveling the mechanics of thermal stress

1209 weathering: rate-effects, size-effects, and scaling laws. J Geophys Res Planets 124(12):3304–3328 Ruesch O, Sefton-Nash E, Vago JL, Küppers M, Pasckert JH, Krohn K, Otto K (2020) In situ fragmentation of lunar blocks and implications for impacts and solarinduced thermal stresses. Icarus 336:113431 Simmons G, Cooper HW (1978, August) Thermal cycling cracks in three igneous rocks. In: International journal of rock mechanics and mining sciences & geomechanics abstracts, vol 15, No. 4, Pergamon, pp 145–148 Thirumalai K, Demou SG (1970) Effect of reduced pressure on thermal-expansion behavior of rocks and its significance to thermal fragmentation. J Appl Phys 41(13):5147–5151 Turcotte DL, Schubert G (2002) Geodynamics. Cambridge University Press Ugural AC, Fenster SK (2003) Advanced strength and applied elasticity. Pearson Education Viles H, Ehlmann B, Wilson CF, Cebula T, Page M, Bourke M (2010) Simulating weathering of basalt on Mars and Earth by thermal cycling. Geophys Res Lett 37(18) Warren K, Eppes MC, Swami S, Garbini J, Putkonen J (2013) Automated field detection of rock fracturing, microclimate, and diurnal rock temperature and strain fields. Geosci Instrum Methods Data Syst 2(2): 275–288

Thermophysical Behavior of the Lunar Surface K. Durga Prasad Planetary Sciences Division, Physical Research Laboratory, Ahmedabad, India

Introduction Moon is a favorable target for exploration not just because of its accessibility but also because of the reason that its pristine surface hides behind several key secrets related to the formation and evolution of our solar system. Knowledge of heat flow on the Moon provides us an insight into its geophysical characterization and thermal evolution (Jaeger 1959; Krotikov and Troitskiĭ 1964; Langseth et al. 1976; Warren and Rasmussen 1987; Wieczorek and Huang 2006; Wieczorek and Phillips 2000; Jolliff et al. 2000; Lawrence et al. 2000). The evolution of planets, the

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terrestrial ones in particular, is thermodynamically controlled and its present state largely depends upon its initial temperature, size, and thermophysical properties of the material from which it was formed (Breuer and Moore 2007). This indicates that almost any process associated with a planetary body can be tracked to its heat budget. Therefore, the total heat budget of a planetary body plays a vital role in every stage of its evolution, either it could be initial accretion or the geochronology or history of its rocks (Hagermann 2005). Thus, the understanding of the current state of a planetary body directs towards a specific thermal evolution scenario which can be better deciphered by investigating its surface/subsurface thermophysical behavior and interior heat flow at present. Observations from recent missions have completely changed the existing perception about the lunar heat flow and thermophysical behavior. Deep Moon quakes (>500 km) (Kawamura et al. 2017), observations of surface shrinkage (Graben) (Watters et al. 2010), reports about young volcanism (Srivastava et al. 2013), dry granular flows, and mass movements point towards the possible existence of accumulated internal heat. Also, significant to mention at this point are the recent observations of distinct thermophysical behavior of morphologically distinct lunar terranes with significant variations in lunar day-night temperatures (Bandfield et al. 2012; Bauch et al. 2011; Vasavada et al. 2012). In situ measurements of lunar thermophysical behavior and heat flow are possible by placing subsurface heat flow probes. But, lunar subsurface heat flow measurements are often perturbed by external thermal forcing due to solar insolation. In order to have an unperturbed heat flow measurement, the probes have to be placed beyond the thermal skin depth of the longest cycle (>2 m) (Durga Prasad et al. 2016). If the global heat flow of the Moon is planned to be estimated in future either by in situ investigations or through remote sensing observations or numerical modeling, precise estimation of equilibrium boundary between external and internal heat fluxes is necessary to infer the net heat flow on the Moon. This requires an in-depth understanding of the lunar near-

Thermophysical Behavior of the Lunar Surface

surface thermophysical behavior. However, the understanding of lunar surface thermophysical behavior is not straight forward and has several interdependencies and therefore a systematic and numerous detailed in-depth investigations. Even after numerous efforts carried out for several decades, very little is known and thus necessitates the planning of a series of geophysical experiments on the Moon in near future. These may include investigations through orbiters, deployment of thermal probes for in situ studies, and network missions.

Significance of Thermophysical Behavior and Heat Flow on the Moon The most important aspect that makes the Moon unique is its size. Moon is large enough to be termed as a planetary-sized body so that it can be considered as a significant sample of the inner solar system. On the other hand, it is small enough that during its history of 4.56 billion years, there is a reason to believe that it has lost majority of its initial heat. Therefore, it can be thought that the present heat flux is primarily due to the heat generated from radioactive decay of elements in its interior particularly up to a depth of 300 km (Keihm and Langseth 1977). The results obtained from petrological and geochemical analysis done on lunar return samples constrain that a radial differentiation occurred during the early epoch of lunar evolution. This differentiation process purged the heat-generating radioactive isotopes 238 U, 235U, 40K, and 232Th from the interior making them to concentrate in the crust/outer layer of the Moon. Thus, the surface heat flow can be considered to be primarily due to this partitioning of heat-producing elements and therefore the surface heat flow can be related to the total abundance of these radiogenic elements: Presence of these incompatible radioactive elements provided a valuable chemical constraint on the Moon’s bulk composition. These elements remained in liquid state up to the last stage of lunar magma crystallization process and can be treated as a proxy to trace back original magma composition and its mixing efficiency during early crystallization

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processes. The mapping of radioactive elements can reveal information on urKREEP layer between the crust and mantle (Warren and Wasson 1979; Jolliff et al. 2000). Detailed heat flow measurements on the Moon are key to validate the giant-impact hypothesis and provide possible reasons for the heterogeneous crustal thickness and structure. Most of the earlier attempts made were mainly motivated either to understand the surface thermal behavior or the subsurface heat flow. These primarily include telescopic observations and orbiter measurements. In situ data was available only from the two Apollo measurements. However, an in-depth knowledge of the thermophysical behavior of the surface and subsurface is necessary for correct interpretation of the results obtained from such in situ heat flow measurements. Furthermore, these surface/subsurface measurements are also often perturbed by the thermal forcing due to solar insolation. Therefore, systematic understanding of the global thermophysical behavior and its implications on heat loss from lunar interior is needed to estimate the net heat flow of the Moon.

(Wieczorek and Huang 2006). Even arguments in terms of abundance of near-surface radioactive elements and nature of geologic provinces of Apollo 15 and Apollo 17 could not provide an unambiguous explanation (Wieczorek and Phillips 2000; Jolliff et al. 2000; Lawrence et al. 2000). Topographic and stratigraphic perturbations were also observed in the measured heat flow values (Langseth et al. 1976). Hence, Apollo heat flow values do not represent the values for the global Moon geophysical characterization and the regional fluxes are highly variable depending on a number of parameters, viz., local temperatures, topography, composition, nature of subsurface, and so on. Most of the numerical studies of the lunar heat flow were also carried out around three decades ago and hence cannot be considered to be valid for representation of global scenario and thus require detailed investigation. Details of the earlier measurements and model studies carried out are briefly given below.

Previous Measurements and Inferences

Ground-Based Observations

The present knowledge about the lunar surface thermophysical properties comes from groundbased observations, remote sensing, and laboratory experiments conducted on samples returned by Apollo missions. Numerical models, mostly one dimensional in nature, combined with Apollo sample return measurements have also aided in inferring information about surface and subsurface temperatures and their variations. However, the major part of the present understanding of lunar heat flow and thermophysical behavior comes from mathematical modeling and analyses of in situ data from Apollo missions (Jaeger 1959; Krotikov and Troitskiĭ 1964). Although analyses of Apollo data placed constraints on the heat flow of the Moon (Langseth et al. 1976; Keihm and Langseth 1973), clear discrepancies exist (Langseth et al. 1976; Warren and Rasmussen 1987) which lead to the reanalysis of data

Moon’s thermal emission, particularly in infrared wavelengths, was being observed by means of Earth-based telescopic observatories since several decades (Mendell and Low 1975; Pettit and Nicholson 1930; Saari and Shorthill 1972; Sinton 1962). These were the first investigations that have provided the first cut idea about various thermophysical properties of the lunar surface. According to these measurements, it was inferred that albedo is the primary factor dictating the daytime surface temperatures of the Moon. Other important factors that affect surface temperatures are surface roughness, local topography, and latitude. According to Pettit and Nicholson (1930), the lunar surface emission was non-Lambertian. Petit and Nicholson provided the general form of temperature variation at lunar equator as shown in Fig. 1. They have suggested that at equator, the surface temperature follows a cosine function of the incidence angle cos1/6(i) against the expected cos1/4(i) for a

Remote Observations The efforts made so far to understand the lunar thermophysical behavior through Earth-based telescopic observations and orbiter observations.

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Thermophysical Behavior of the Lunar Surface 400 1939 NEAR CENTER

1 HOUR OF TIME

1927 NEAR

300 SOUTH RIM

TEMP. °K

TOTALITY

200

PARTIAL PHASE

PARTIAL PHASE

140 TIME

Thermophysical Behavior of the Lunar Surface, Fig. 1 Eclipse cooling curve near south rim and center of the Moon obtained on 24 June 1927 and 27 Oct 1939 by

ground-based observations (from Saari and Shorthill 1966, https://apps.dtic.mil/dtic/tr/fulltext/u2/645548.pdf)

Lambertian source. Sinton (1962) and Pettit and Nicholson (1930), both measured the thermal emission from the lunar subsolar point as a function of phase angle. Pettit and Nicholson found that the surface temperature decreases with increasing phase angle. Sinton (1962) found higher temperature than expected from a Lambertian surface from subsolar point on a full Moon. Saari and Shorthill (1972) predicted the limbs of the full Moon to be even warmer than that suggested by Pettit and Nicholson (1930) proposing a relationship for full Moon brightness temperature (TB), as a function of angular distance (i) from the subsolar point as given in Eq. (1).

measurements were used. Accordingly, temperatures for sunrise were estimated to be ~109 K, while for sunset, these values were found to be between 170 and 181 K. The values for noon time temperatures were measured between 374 and 393 K.

TB ¼ 324:2 þ 72:6 cosðiÞ

ð1Þ

However, these anomalies have been attributed to surface roughness. Nighttime temperatures were limited from these observations and temperatures less than 100 K were difficult to be determined. Daytime measurements were further associated with the difficulty of separating emitted and reflected radiation. Therefore, eclipse

Orbiter Observations

Two lunar orbiters, Clementine and Lunar Reconnaissance Orbiter, carried instruments capable of measuring thermal emission from lunar surface to derive lunar surface thermophysical characteristics. Observations from the long-wavelength infrared camera onboard Clementine Orbiter provided global mosaic of daytime temperatures near local noon with a resolution of 1o/pixel. The noon time surface temperatures were predicted to vary around ~200 K between equator and poles. Mare regions were estimated to be ~10 K warmer than the highlands which were attributed to albedo than to thermal inertia or composition of the surface. The Diviner Lunar Radiometer Experiment (DLRE) LRO has been carrying out infrared probing of the lunar surface in four thermal channels operating in the range from 12 to 400 mm that

Thermophysical Behavior of the Lunar Surface

covers the complete temperature range expected on the lunar surface (Paige et al. 2010; Vasavada et al. 2012). Global maps for daytime and nighttime lunar surface temperatures at a resolution of 0.5 /pixel were generated from Diviner observations (Williams et al. 2017). The Diviner results shown in Fig. 2 indicate an average maximum and minimum temperatures of 392.3 and 94.3 K, respectively, for the equator, while the day-night mean equatorial surface temperature was found to be 392.3 K thus exhibiting a variation of ~300 K during an entire lunar day as predicted earlier. The regions near south pole appeared to be ~11 K

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warmer than that of the north pole. As a result of delay between heating and cooling of the regolith due to thermal inertia, morning and afternoon temperatures exhibited an asymmetrical distribution. Albedo and surface roughness have been found to play a dominant role in deriving the surface temperatures. LRO Orbiting at an altitude of 50 km, global and local temperature maps were possible at a spatial resolution of 320  160 m. Several thermal anomalous features have been identified from nighttime temperature observations. These anomalies which are found from thermal inertia maps are seen to be associated

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Thermophysical Behavior of the Lunar Surface, Fig. 2 Diviner derived lunar global surface bolometric temperatures. (Image source: Williams et al. 2017)

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Thermophysical Behavior of the Lunar Surface

In Situ Observations The first-ever in situ measurement of surface temperature and thermophysical characteristics was

accomplished by Surveyor-I spacecraft in 1966 (Lucas et al. 1967). Although not a direct measurement, the brightness temperature in the vicinity of the lander observed by the outside cabin of the spacecraft was used to derive the lunar surface temperature. The surveyor spacecraft depicting the two cabins used for inferring surface temperatures and the retrieved temperature profiles are shown in Figs. 3 and 4, respectively. Apollo missions have provided the first opportunity for direct measurement of in situ surface and subsurface temperatures. Apollo missions have used heat flow probes which were emplaced in drilled boreholes. Several temperature measurements were

Thermophysical Behavior of the Lunar Surface, Fig. 3 Surveyor spacecraft configuration – the two surface electronics compartment where the TB measurements

were done is highlighted (Image source: Wikimedia, Modfied after https://upload.wikimedia.org/wikipedia/ commons/2/2a/Surveyor_NASA_lunar_lander.jpg)

with features like craters, fresh impacts, and rock abundances (Bandfield et al. 2011). Using the microwave brightness temperature measurements from microwave radiometer onboard Chang’E-1 and Chang’E-2 missions, the surface temperatures have been estimated. Since microwave penetrates the subsurface, subsurface temperatures were also estimated and an empirical relationship was provided for latitudinal variation of lunar surface and subsurface temperatures.

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Thermophysical Behavior of the Lunar Surface, Fig. 4 First-ever in situ lunar surface temperature profile measured by Surveyor 1 spacecraft (Lucas et al. 1967)

done during Apollo missions. Apollo 15 and Apollo 17 could successfully measure lunar surface and subsurface temperatures and their diurnal and annual variability. The calculated lunar surface temperatures for all the four probes of Apollo 15 and Apollo 17 Heat Flow Experiments (HFEs) are shown in Fig. 5. Figure 6 shows the profiles of thermal conductivity and density deemed to be fit based on the cool down histories and thermal model for the surface temperature profiles. For Apollo 17 landing site, a variation of more than 280 K has been observed between the maximum and minimum temperatures of ~384 and ~102 K, respectively. Apollo 17 site reported nearly 10 K higher temperatures throughout the night. Rapid variation in temperature during sunrise and sunset was seen due to absence of atmosphere on the Moon. The midnight temperatures of ~106 K obtained from Apollo showed a good agreement with previous infrared observations of Saari (1964). The Apollo missions also carried an Infrared Scanning Radiometer (ISR) on the command service module in addition to the in situ probes. The ISR which operated between 1.2 and 70 mm (Mendell and Low 1974, 1975) provided thermal

maps for the surface in the vicinity of the landing modules. Thermal anomalies were found in ISR measurements which were explained by the presence of impact craters (Mendell 1976). In addition to surface and subsurface temperature measurements, the deployed probes at Apollo site provide heat flow measurements. Whatever the present understanding we have about the lunar heat flow comes from the only two in situ measurements carried out during Apollo 15 and Apollo 17 missions. Surface and subsurface temperature measurements along with measured thermal conductivity values at Apollo 15 and Apollo 17 sites were used to derive heat flow values for these sites. Although analysis of Apollo data placed constraints on net heat flow of the Moon (Wieczorek and Huang 2006; Saito et al. 2008), certain discrepancies still exist. Figure 7 presents the summary of thermal gradient and temperature profile obtained from Apollo 15 and Apollo 17 heat flow probes. Values of 21 and 16 mW/m2 have been derived from Apollo 15 and Apollo 17 in situ heat flow measurements, respectively (Langseth et al. 1972a, b, 1973, 1976). These measurements were carried out at

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temperature drift also existed which was actually attributed to astronaut-related disturbances (Langseth et al. 1976). However, this could be thought to be a result of even long-term disturbances due to 18.6-year precision cycle (Wieczorek and Huang 2006; Saito et al. 2008). Local topography can also have significant effect on the heat flow values which was not considered in those calculations (Wieczorek and Huang 2006; Saito et al. 2008). Therefore, no clear conclusion could be made from Apollo measurements due to various associated uncertainties. Moreover, Apollo in situ measurements were site-specific and profiled temperatures at depths beyond ~30–2.5 m. Near-surface behavior was inferred only from extrapolation.

Thermophysical Behavior of the Lunar Surface, Fig. 5 The calculated lunar surface temperature during a lunation using the cable thermocouple data from Apollo 15 and Apollo 17 HFEs (Modified after Langseth and Keihm 1974, Image Source for Inset picture: https:// www.hq.nasa.gov/alsj/HamishALSEP.html)

Hadley Rille and Taurus-Littrow valley. However, there were uncertainties of around
15%. This variation has been considered due to the nature of the underlying regolith’s thermal conductivity. But, these are the only measurements available till date which even serve today as the basis for the estimation of the present thermal state of the Moon. However, independent seismic- and magnetic-based measurements also exist (Pullan and Lambeck 1980; Hood et al. 1982). Rasmussen and Warren (1985) and Warren and Rasmussen (1987) have considered and corrected for local heat flow effects which provided a global average for heat flow as 12 mW/m2. A bulk uranium content of 20–21 ppb was estimated from these measurements. It was thought that several other perturbations such as long-term subsurface

Model Studies Several models have been developed during the past 60 years to numerically estimate the lunar surface and subsurface temperatures. The basic purpose of these models was to interpret observations obtained either from Earth-based telescopes or orbiters. Most of these models were used to predict lunar surface temperatures as a function of various parameters, viz., solar illumination, dielectric constants, emissivity, albedo, and internal thermal flux of the lunar regolith. Because of the lack of sufficient and systematic experimental observations, the model results have been widely accepted and used till now. Wesselink (1948) developed a thermal model to explain the variations observed in the lunar surface temperature particularly during eclipse and lunation. The model was developed by considering that the temperature of the surface follows a cosine function or Fourier series. Using this model, the thermal conductivity and the nature of lunar surface was estimated. The model was also used to determine the heat influx within the first few millimeters of the regolith due to the surface temperature variations. A lunar surface temperature model was developed by Jaeger (1953) where he considered Moon as a homogeneous body with a thin extremely low conducting layer over a relatively highly conducting subsurface. However, the model could not unambiguously explain the

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Thermophysical Behavior of the Lunar Surface, Fig. 6 Model-fit thermal conductivity profiles using the cable thermocouple data from Apollo 15 and Apollo 17 HFEs (From Langseth and Keihm 1974)

CONDUCTIVITY Wx10–4/cm-K

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Thermophysical Behavior of the Lunar Surface, Fig. 7 Temperature-depth curves for four lunar heat flow measurements (From Langseth and Keihm 1974)

experimental observations. A series of models for lunar surface temperatures were proposed by Linsky (1966). The models include both temperature-dependent and temperature-

independent properties to derive surface temperatures of the Moon. The results from the model were in agreement with observations on a gross scale; however, there were some discrepancies

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observed as well. The temperatures and thermophysical properties of the lunar outermost layer were simulated by Jones et al. (1975) by considering parameter interdependency like density- and temperature-dependent thermal conductivity, temperature-dependent specific heat, and so on. The surface temperatures calculated from the model were in good agreement with the infrared and millimeter wave observational data obtained from Earth-based observatories. The model could also reproduce the known properties of the lunar outermost layer known at that time. Jones et al. (1975) also estimated that the density of the outermost layer could vary between 0.7 and 0.9 g/cc while that of the layer underneath could be ~2 g/cc. Based on the thermal conductivity equation and principle of conservation of energy, Racca (1995) developed two different surface models to derive a relationship between the surface temperature and the solar illumination, thermal emission, absorption and the internal thermal flux in steady state, and transient behavior. The significant improvement of this model is the consideration of internal thermal flux into consideration apart from other parameters. Vasavada et al. (1999) developed a one-dimensional thermal model based on model developed by Mitchell and de Pater (1994) for interpreting thermal emission of Mercury. They have developed this model primarily to explain the occurrence and stability of ice deposits in the polar regions of Moon and Mercury. They have considered a two-layer model with 2 cm thick top insulating layer of bulk density 1.3 g/cc followed by a dense conductive layer having bulk density of 1.8 g/cc. This model also accounts for temperature-dependent thermal conductivity. Using the model combined with surface topography, they derived the surface temperatures of flat surface and lunar cold traps. Further modifications in the model proposed by Vasavada et al. (1999) were applied to validate and explain LRO Diviner Radiometer observations (Vasavada et al. 2012). The modifications applied were based on considering solar irradiance of 1360.8 W/m2 for minimum conditions and ephemeris appropriate during the perihelion of Diviner observations. Using the thermal

Thermophysical Behavior of the Lunar Surface

conductivity equation, and knowledge of thermophysical properties obtained from the returned samples, some researchers have derived surface temperatures of the Moon which takes into account an improved transient temperature model that includes time-dependent solar irradiance. The results were validated using ground truth obtained from Apollo 15 and Apollo 17 missions. Miyahara et al. (2008) developed a lunar surface temperature model to reconstruct the historical solar irradiance profiles from Apollo 15 and Apollo 17 lunar borehole temperatures. This model also solves one-dimensional heat flow equation along with thermal and radiative properties inferred from Apollo measurements and ephemeris from JPL. Using the model, Miyahara et al. (2008) computed the present-day lunar regolith temperatures for tropical, mid-latitude, and polar regions of the Moon. However, certain important constraints such as terrain, bedrock, internal heat variations, etc. which can affect the borehole temperatures were not considered. Zhiguo et al. (2014) used an improved Racca model to estimate the effect of topography on lunar surface temperatures. The topography input for the model is taken from Chang’E-1 topography data. The results showed that there is a significant dependence of topography on the surface temperature variations. Topographic effect can provide variations as large as 150 K in certain cases. Durga Prasad et al. (2015) recently proposed a new comprehensive three-dimensional finite element thermophysical model to address the local to regional scale variations in the thermophysical behavior of the Moon. This model takes into account the complex topography and parametric-based variation in the physical properties of lunar surface.

Status of Current Understanding As detailed in the earlier section, the current state of knowledge about lunar surface thermophysical behavior comes from ground-based observations, orbiter observations, and laboratory experiments conducted on samples returned

Thermophysical Behavior of the Lunar Surface

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Thermophysical Behavior of the Lunar Surface, Fig. 8 Schematic representation of two-layer nature of lunar surface from present understanding (From Durga Prasad et al., 2015)

by Apollo missions. The remote sensing observations when integrated with numerical models, mostly one dimensional in nature, and combined with Apollo sample return measurements have provided information about surface and subsurface temperatures and their variations. The only contribution of the in situ measurements of lunar heat flow is from the two landing missions: Apollo 15 and Apollo 17. From Apollo measurements, Keihm and Langseth (1977) estimated a mean heat flow of 14–18 mW/m2 on the basis of Apollo 15 and Apollo 17 heat flow values of 21 and 16 mW/m2, respectively (Langseth et al. 1976). This range of heat flow values corresponds approximately to a global lunar U concentration of 33–44 ppb, if a steady-state balance between heat loss (flow) and heat production is assumed. The estimated bulk uranium concentration of Moon is much higher than that of Earth’s mantle. However, derivation of a globally representative average of the heat flow from measurement at the two Apollo sites is necessarily difficult. The bulk abundance of refractory elements on the Moon, as compared with that in the Earth, is important for delineating the lunar origin. Therefore, more measurements of heat flow at different geological settings are desirable to derive a reliable global

heat flow average and hence a global uranium abundance on the moon. Based on the present understanding, the lunar surface can be best thought to be represented by a two-layer model having two distinct thermal regimes as shown in Fig. 8 (Durga Prasad et al. 2015). It consists of an uppermost porous layer (~2 cm thick) of extremely low thermal conductivity followed by a regolith layer beneath with relatively higher thermal conductivity. The surface is subjected to solar insolation during lunar day-night conditions. The temperature and heat flux at the surface of the Moon are determined mainly by the solar energy impinging on the surface during 29.5-day lunation cycle. However, the extremely low conductive nature of the lunar regolith acts a thermal blanket and strongly inhibits the flow of external heat into and out of the subsurface which is attenuated to nearly undetectable amplitude at a depth of ~50 cm. However, annual thermal wave effects are expected to manifest till depths of nearly few meters. Beyond these depths, the thermal regime is dominated by heat flow from the lunar interior. Precise estimation of this equilibrium boundary between external and internal heat fluxes is therefore necessary to infer the net heat flow on the Moon. However, the extent and spatial

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distribution of this regime is not well constrained. Due to lack of any atmospheric or tectonic processes, interpretation of lunar heat flow has been thought to be relatively less complex under the assumption of thermal steady state. But the thermophysical behavior of the upper most (~1 m) surficial layers of the moon is highly complex because of their complex geometric, thermal, and radiative properties. Therefore, thermal exchange due to solar insolation within these layers becomes more complex exhibiting a kind of non-steady-state behavior. In particular, the surficial ~10–20 cm strata with an outermost porous layer (~ few cms) is expected to exhibit significant parametric, compositional, stratigraphical, and latitudinal variations. Although the presence of ~2 cm outermost porous layer on the Moon has been validated through remote sensing and numerical modeling, its nature, exact thickness, and spatial extent are not tightly constrained due to lack of sufficient experimental data which demands for attempts of in situ measurements. Since this porous layer principally dictates the propagation of solar heat influx to the interior layers, an in-depth understanding of the thermophysical behavior of lunar surficial layer and its stratigraphic and spatial variation is needed for a better estimation of the equilibrium boundary due to external and internal heat fluxes. Furthermore, thermal variations within the regolith due to lunar day-night cycle and 18.6-year precession of the lunar orbit also constrain the determination of regolith boundary for thermal exchange due to internal and external heat sources. The lunar surface, in particular the upper most porous layer, exhibits a complex thermophysical behavior and has not been well understood so far. The principal reason for this is the interdependence of various physical properties and their effect on heat exchange within the medium and the exotic temperature and vacuum conditions of the Moon. Due to the large daynight temperature gradients of ~300 K on the Moon, the basic thermophysical properties, viz., thermal conductivity and specific heat exhibit significant variations. Further, the vacuum conditions make the story more complicated. Direct

Thermophysical Behavior of the Lunar Surface

investigations of these properties for the surface of the Moon are available from samples returned from Apollo missions, but these samples are limited and mostly biased to particular regions and conditions of sampling. At this point, the only way one can improve the current understanding about these aspects is through laboratory experiments and numerical simulations. Lunar Surface Thermophysical Properties and Their Interdependence The thermophysical properties of a material are defined as those properties which alter with temperature but do not change due to the chemical nature of the material. These properties basically play an important role in the storage and transfer of heat in the medium and are primarily dependent on temperature, pressure, and composition. The relevant thermophysical properties of lunar surface are thermal conductivity, thermal diffusivity, and specific heat. Since heat flow is a function of thermal conductivity and temperature gradient of a medium, heat flow is also used as an indicator of the thermophysical behavior of the medium. The thermal response of the lunar surface to solar insolation has a direct dependence on the surface bulk density, specific heat, and thermal conductivity. Thermal conductivity is the dominant parameter that dictates thermal transport within the lunar soil. In situ measurements of subsurface thermal conductivity were measured at two Apollo locations. The results indicate that the thermal conductivity increases significantly with depth. Results from these measurements also concluded that the top few cm (1–2 cm) exhibits an extremely low thermal conductivities of the order of 1.5  105 W/cm K while the conductivity increases by five to seven times at depths beyond 2 cm. The thermal conductivity values from two in situ Apollo measurements are summarized in Table 1. These thermal conductivity values show consistency with the density values obtained from the drill cores. Although certain estimates for thermal conductivity are available, they are not well constrained. However, the current understanding is that the lunar surface thermal conductivity for solid, granular regolith

Thermophysical Behavior of the Lunar Surface Thermophysical Behavior of the Lunar Surface, Table 1 Summary of in situ thermal conductivity measurements by Apollo 15 and Apollo 17 HFEs (Langseth and Keihm 1974) Thermal conductivity of surface layer Thermal conductivity below 10 cm Mean temperature difference (upper few cm) Mean vertical temperature gradient (subsurface)

1.2–1.5 mW/m. K 10–15 mW/m. K 30–35 K 1.85–1.35 K/m

increases with temperature as a result of two primary heat transfer mechanisms, conduction and radiation. Based on this understanding combined with observational result, a model fit has been obtained for thermal conductivity. However, the thermal conductivity of the outermost fluff is still poorly constrained. Furthermore, the thermal conductivity itself is a function of several parameters, viz., grain size, bulk density/porosity, temperature, and presence of rocks. It was shown from laboratory experiments on analogous samples that the parameters such as grain size, porosity, and vacuum conditions significantly affect the heat flux within the lunar soil. Temporal variability of heat flux within the soil as a function of these parameters from one of the laboratory experiments is shown in Fig. 9 (Durga Prasad and Murty 2014; Durga Prasad et al. 2016). The specific heat capacities of returned sample from several Apollo missions have been measured in laboratory. The specific heat capacity exhibits a strong dependence on temperature and composition. A linear fit has been obtained from various observations for estimating specific heat of the lunar soil. Specific heat also shows a variation with respect to temperature. The experimental specific heat values for Apollo lunar soil sample 15301, 20 from Hadley Apennine base is shown in Table 2 as a function of temperature. Table 3 shows typical value of thermal conductivity and specific heat as a function of temperature for lunar soil having density 1.3 g/ cc. It was shown that the two properties, thermal conductivity and specific heat, do not show significant variation with respect to composition for

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different types of silicates such as glass or basalts (Wechsler et al. 1972; Fountain and West 1970; Birkebak et al. 1970; Winter and Saari 1969; Wechsler and Glaser 1965). Therefore, the thermophysical behavior can be thought to be more dependent on the thermal conductivity and specific heat and relatively less dependent on composition. The parameters such as bulk density and grain size particularly for powdered materials make things more difficult. For example, the bulk density of a solid rock for two different types of silicates may not vary much but the same in the case of powders show significant variation. In this case, porosity plays a dominant role in dictating the thermophysics. Another aspect that can predominantly affect the thermophysical property is the distribution, arrangement, and compaction of the particles as they can substantially change the porosity with an imperceptible change in density. Considering a vacuum environment and porous material, the behavior becomes completely different. Under vacuum environments of airless bodies like Moon, intergranular convection is absent and the bulk thermal conductivity or heat flow is only through particle to particle conduction and interparticle radiation through voids. Therefore, for porous media under vacuum conditions, the effective thermal conductivity is expressed as a sum of solid conduction and interparticle radiation as shown in Eq. (1). leff ¼ ls þ lr

ð2Þ

where “leff” is effective thermal conductivity, “ls” is solid conductivity, and “lr” is radiative transfer. Further, the radiation is highly temperature dependent and its effect is insignificant at low temperatures and dominant at high temperatures. For huge diurnal temperature excursions on the Moon, the nighttime heat transfer can be thought to be only due to conduction while during day time it is contribution from both conduction and radiation. Although it was estimated that the thermal conductivity of lunar soils during lunar day can vary by a factor of 3 (Urquhart and Jakosky 1997), unique trend or behavior is not known. In

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Thermophysical Behavior of the Lunar Surface, Fig. 9 Experimental result showing temporal variation of heat flow within the sample stratigraphy for variable pressures and grain sizes (From Durga Prasad and Murty, 2014)

Thermophysical Behavior of the Lunar Surface, Table 2 Laboratory measured specific heat values from Hadley Apennine base sample 15301, 20 (Hemingway et al. 1973) Temperature, K 83.98 95.46 104.47 114.25 125.20 135.82 146.20 156.58 167.16 177.64 188.03

Specific heat, J/(g  K) 0.2218 0.2569 0.2870 0.3197 0.3498 0.3824 0.4171 0.4443 0.4724 0.5004 0.5280

Temperature, K 198.34 208.28 217.89 227.13 237.00 247.16 257.76 268.32 278.68 289.07 –

addition to all these effects, specific heat shows significant change at higher temperatures and must be accounted for when inferring thermophysical behavior of the surface of Moon. Also, the heat flow within the regolith is dependent strongly up on the stratigraphy of the regolith.

Specific heat, J/(g  K) 0.5540 0.5774 0.6004 0.6205 0.6427 0.6636 0.6862 0.7063 0.7276 0.7481 –

Temperature, K 299.50 303.96 310.91 306.01 315.65 325.09 334.25 343.42 352.89 363.10 –

Specific heat, J/(g  K) 0.7669 0.7778 0.7847 0.7782 0.7958 0.8125 0.8280 0.8452 0.8619 0.8765 –

This is also perturbed by the presence of rocks and variations in local topography. Based on these discussions, it is clearly evident that the thermophysical behavior of lunar surface is very much complicated and several systematic studies and in situ measurements are needed in future.

Thermophysical Behavior of the Lunar Surface

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Thermophysical Behavior of the Lunar Surface, Table 3 Thermal parameters and their variation with temperature measured from Apollo samples (Hemingway et al. 1973) Temperature,  K 100 150 250 300 350

Thermal conductivity k, W m1 K1 0.0007 0.0008 0.0011 0.0014 0.0017

Specific heat Cp, J kg1 K1 275.7 433.9 672.4 758.1 848.9

Open Questions There are several open questions that need to be addressed to have a comprehensive understanding of the global lunar heat flow/thermophysical behavior. Some of them, although not in the scope of this work, are listed below. (i) How do the morphology, petrology, and stratigraphy of the Moon influence the heat flow and thermophysical characteristics? (ii) How do the surface and subsurface thermal properties and heat flow vary with depth? (iii) What is the regolith boundary for thermal exchange due to internal and external heat to constrain the abundance and distribution of radioactive elements in lunar interior? (iv) What is the thermophysical behavior of nearsurface layers due to lunation and annual heat wave attenuation? How does it vary with location? (v) How do the thermophysical parameters behave in the case of megaregolith insulation (Warren and Rasmussen 1987) which has been completely neglected so far?

Summary and Future Course of Directions Systematic measurements of heat loss from lunar interior and its spatial variation are needed to accurately quantify the net heat flow on the Moon. We have only two in situ measurements from Apollo missions which are not sufficient to understand the global thermophysical behavior and heat flow of the Moon. Although heat flow

Thermal parameter γ, M2 s1/2 KJ1 0.06313 0.04707 0.03225 0.02692 0.02309

measurements will be of top priority for future in situ geophysical exploration of the Moon, no such mission is in the offing at least for the next decade. Recently, Chandra’s Surface Thermophysical Experiment (ChaSTE) was flown onboard Chandrayaan-2 Lander, Vikram, to investigate the surface temperatures and thermophysical behavior at a high latitude lunar site (https:// www.isro.gov.in/chandrayaan2-payloads). Globally, no other such effort seems to be planned in near future. The lunar surface, in particular the upper most porous layer, exhibits a complex thermophysical behavior and has not been well understood so far. Lunar surface and subsurface temperatures of the Moon are dictated by a complex interplay of number of parameters. These surface and subsurface temperatures also manifest significant latitudinal, stratigraphic, and topographical variability. Due to the large day-night temperature gradients of ~300 K on the Moon, the basic thermophysical properties, viz., thermal conductivity and specific heat also exhibit significant variations. Further, the vacuum conditions make the story more complicated. Also, the heat flow within the regolith is dependent strongly up on the stratigraphy of the regolith. This is also perturbed by the presence of rocks and variations in local topography. Thermal conductivity and specific heat show less variation with composition between different types of silicates, such as basalt or glass but vary largely with rock or soil (Wechsler et al. 1972; Birkebak et al. 1970; Winter and Saari 1969). Therefore, the composition of the sample materials is less constrained. On the other hand, heat capacity varies strongly with temperature (by a factor of

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4 within 80–390 K) and a constant value cannot be used. Apollo Heat Flow Experiments were restricted to only equatorial latitudes. Thermal behavior at mid and particularly at high latitudes is not at all known. One way of understanding these aspects before any future in situ measurements are available is to conduct laboratory experiments on analogous sample under simulated lunar environment. Although some efforts are being made (Durga Prasad et al. 2017), extensive studies are needed. However, laboratory experiments have certain limitations in terms of simulating parametric variation and long-term variability. This can be overcome by augmenting the laboratory experiments with comprehensive numerical simulations. Regional geophysical modeling of Apollo sites was also attempted using a three-dimensional model but on a global scale (Siegler and Smrekar 2014). Still, a comprehensive model addressing the local to regional scale variations, particularly considering the role of uppermost fluffy layer on lunar thermophysical behavior, is not available. Apart from these, the most prospective thing to accomplish in future would be to plan for a series of geophysical measurements through international cooperation, possibly through establishing an international lunar network.

Cross-References ▶ Lunar Surface, Bulk Density and Porosity ▶ Regolith Physical Properties ▶ Surface and Near-Surface Thermal Environment of the Moon

References Bandfield JL, Ghent RR, Vasavada AR, Paige DA, Lawrence SJ, Robinson MS (2011) Lunar surface rock abundance and regolith fine temperatures derived from LRO Diviner Radiometer data. J Geophys Res 116:E00H02 Bandfield JL, Song E, Hayne PO, Ghent RR, Paige DA (2012) Lunar “cold spots”: a new class of thermophysically and morphologically distinct craters. In: Proceedings of 43rd lunar and planetary science conference – 2012, #1487

Thermophysical Behavior of the Lunar Surface Bauch KE, Hiesinger H, Robinson MS, Scholten F (2011) Thermophysical properties of selected lunar study regions determined from LROC and Diviner data. In: Proceedings of 42nd lunar and planetary science conference – 2011, #2278 Birkebak RC, Cremers CJ, Dawson JP (1970) Thermal radiation properties and thermal conductivity of lunar materials. Science 167:724–726 Breuer D, Moore WB (2007) Dynamics and thermal history of the terrestrial planets, Moon and Io. Treat Geophys 10:299–348 Durga Prasad K, Murty SVS (2014) Effect of grain size and porosity on surface heat flux on the Moon. In: Proceedings of 45th lunar and planetary science conference – 2014, #1236 Durga Prasad K, Rai VK, Murty SVS (2015) A thermal model to study the effect of top porous layer on subsurface heat flow of Moon. In: Proceedings of 46th lunar and planetary science conference – 2015, #1768 Durga Prasad K, Rai VK, Murty SVS (2016) A comprehensive thermal model for an insight into diurnal and latitude variability of lunar subsurface temperatures. In: Proceedings of 47th lunar and planetary science conference – 2016, #1290 Durga Prasad K, Rai VK, Murty SVS (2017) Thermophysical behaviour of lunar analogues under simulated lunar environment. In: Proceedings of 48th lunar and planetary science conference – 2017, #1414 Fountain JA, West EA (1970) Thermal conductivity of particulate basalt as a function of density in simulated lunar and Martian environments. J Geophys Res 75 (20):4063–4069 Hagermann A (2005) Planetary heat flow measurements. Philos Trans R Soc A. https://doi.org/10.1098/rsta.2 005.1664 Hemingway BS, Robie RA, Wilson WH (1973) Specific heats of lunar soils, basalt, and breccias from the Apollo 14, 15, and 16 landing sites, between 90 and 350 K. In Lunar and Planetary Science Conference Proceedings (Vol. 4, p. 2481) Hood LL, Herbert F, Sonett CP (1982) Further efforts to limit lunar internal temperatures from electrical conductivity determinations. J Geophys Res 87:A109– A116 Jaeger JC (1953) The surface temperature of the moon. Aust J Phys 6(1):10–21 Jaeger JC (1959) Sub-surface temperatures on the Moon. Nature 183:1316–1317 Jolliff BL, Gillis JJ, Haskin LA, Korotev RL, Wieczorek MA (2000) Major lunar crustal terranes: surface expressions and crust–mantle origins. J Geophys Res 105(E2):4197–4216 Jones WP, Watkins JR, Calvert TA (1975) Temperatures and thermophysical properties of the lunar outermost layer. Moon 13(4):475–494 Kawamura T, Lognonne P, Nishikawa Y, Tanaka S (2017) Evaluation of deep moonquake source parameters: implication for fault characteristics and thermal state. J Geophys Res 122(7):1487–1504

Thermophysical Behavior of the Lunar Surface Keihm SJ, Langseth Jr, MG (1973) Surface brightness temperatures at the Apollo 17 heat flow site: Thermal conductivity of the upper 15 cm of regolith. In Lunar and Planetary Science Conference Proceedings (Vol. 4, p. 2503) Keihm SJ, Langseth MG (1977) Lunar thermal regime to 300 km. In: Proceedings of 8th lunar science conference, pp 499–514 Krotikov VD, Troitskiĭ VS (1964) Radio emission and nature of the moon. Sov Phys Usp 6(6):841–871 Langseth MG, Clark SP, Chute JL, Keihm SJ, Wechsler AE (1972a) Heat flow experiment, in Apollo 15: preliminary science report, Rep. SP-289, ch 11. National Aeronautics and Space Administration, Washington, DC, pp 1–23 Langseth MG, Clark SP, Chute JL, Keihm SJ, Wechsler AE (1972b) The Apollo 15 lunar heat-flow measurement. Earth Moon Planet 4(3–4):390–410. https://doi.org/10. 1007/BF00562006 Langseth MG, Keihm SJ, Chute JL (1973) Heat flow experiment, in Apollo 17: preliminary science report, Rep. SP-330, ch 9. National Aeronautics and Space Administration, Washington, DC, pp 1–24 Langseth MG, Keihm SJ (1974) In-situ measurements of lunar heat flow. In Soviet-American Conference on the Cosmochemistry of the Moon and Planets Langseth MG, Keihm SJ, Peters K (1976) Revised lunar heat-flow values. In: Proceedings of 7th lunar science conference, vol 7, pp 3143–3171 Lawrence DJ, Feldman WC, Barraclough BL, Binder AB, Elphic RC, Maurice S, Miller MC, Prettyman TH (2000) Thorium abundances on the lunar surface. J Geophys Res 105:20 Linsky JL (1966) Models of the lunar surface including temperature-dependent thermal properties. Icarus 5:606–634 Lucas JW, Conel JE, Hagemeyer WA, Garipay RR, Sari JM (1967) Lunar surface thermal characteristics from Surveyor 1 in JPL Technical Report No. 32-1023 (Surveyor I Mission Report – Part II: Scientific Data and Results). Website: https://ntrs.nasa.gov/archive/ nasa/casi.ntrs.nasa.gov/19670000738.pdf Mendell WW (1976) Degradation of large, period II lunar craters. In: Proceedings of 7th lunar science conference, pp 2705–2717 Mendell WW, Low FJ (1974) Preliminary results of the Apollo 17 infrared scanning radiometer, Moon, 9:97– 103, https://doi.org/10.1007/BF00565396 Mendell WW, Low FJ (1975) Infrared orbital mapping of lunar features. In: Proceedings of 6th lunar and planetary science conference, pp 2711–2719 Mitchell DL, de Pater I (1994) Microwave imaging of Mercury’s thermal emission at wavelengths from 0.3 to 20.5 cm. Icarus 110:2–32 Miyahara H, Wen G, Cahalan RF, Ohmura A (2008) Deriving historical total solar irradiance from lunar borehole temperatures. Geophys Res Lett 35(2) Paige DA, Siegler MA, Zhang JA, Hayne PO, Foote EJ, Bennet KA, Vasavada AR, Greenhagen BT, Schofield

1225 JT, McCleese DJ, Foote MC, DeJong E, Bills BG, Hartford W, Murray BC, Allen CC, Snook K, Soderblom LA, Calcutt S, Taylor FW, Bowles NE, Bandfield JL, Elphic R, Ghent R, Glotch TD, Wyatt MB, Lucey PG (2010) Diviner Lunar Radiometer observations of cold traps in the Moon’s south polar region. Science 330:479 Pettit E, Nicholson SB (1930) Lunar radiation and temperatures. Astrophys J 71:102–135 Pullan S, Lambeck K (1980) On constraining lunar mantle temperatures from gravity data. In: Proceedings of 11th lunar and planetary science conference, vol 3, pp 2031–2041 Racca GD (1995) Moon surface thermal characteristics for moon orbiting spacecraft thermal analysis. Planet Space Sci 43(6):835–842 Rasmussen KL, Warren PH (1985) Megaregolith thickness, heat flow, and the bulk composition of the Moon. Nature 313:121–124. https://doi.org/10.1038/ 313121a0 Saari JM, Shorthill RW (1966) Review of lunar infrared measurements. Technical document, Boeing Scientific Research Laboratories, D1.82.0586 Saari JM, Shorthill RW, Winter DF (1972) The sunlit lunar surface. The moon, 5(1–2):179–199 Saito Y, Tanaka S, Horai K, Hagermann A (2008) The long term temperature variation in the lunar subsurface. Lunar Planet Sci XXXIX:Abstract 1663 Siegler MA, Smrekar SE (2014) Lunar heat flow: regional prospective of the Apollo landing sites. J Geophys Res 119(1):47–63 Sinton WM (1962) In: Kopal Z (ed) Physics and astronomy of the Moon. Academic Press, New York Srivastava N, Kumar D, Gupta RP (2013) Young viscous flows in the Lowell crater of Orientale basin, Moon: impact melts or volcanic eruptions? Planet Space Sci 87:37–45 Urquhart ML, Jakosky BJ (1997) Lunar thermal emission and remote determination of surface properties. J Geophys Res 102:10959–10969 Vasavada AR, Paige DA, Wood SE (1999) Nearsurface temperatures on Mercury and the Moon and the stability of polar ice deposits. Icarus 141(2):179–193 Vasavada AR, Bandfield JL, Greenhagen BT, Hayne PO, Siegler MA, Williams J, Paige DA (2012) Lunar equatorial surface temperatures and regolith properties from the Diviner Lunar Radiometer Experiment. J Geophys Res 117:E00H18. https://doi.org/10.1029/2011JE003987 Warren PH, Rasmussen KL (1987) Megaregolith insulation, internal temperatures, and bulk uranium content of the Moon. J Geophys Res 92:3453–3465. https://doi. org/10.1029/JB092iB05p03453 Warren PH, Wasson JT (1979) The origin of KREEP. Rev Geophys Space Phys 17:73–88 Watters T, Robinson MS, Beyer RA, Banks ME, Bell JF III, Pritchard ME, Hiesenger H, Bogert CH, Thomas PC, Turtle EP, Williams NR (2010) Evidence of recent

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1226 thrust faulting on the Moon revealed by the Lunar Reconnaissance Orbiter Camera. Science 329:936–940 Wechsler AE, Glaser PE (1965) Pressure effects on postulated lunar materials. Icarus 4:335–352 Wechsler AE, Glaser PE, Little AD, Fountain JA (1972) Thermal properties of granulated materials. In: Lucas JW (ed) Thermal characteristics of the Moon. MIT Press, Cambridge, MA. Prog Astronaut Aeronaut 28:51–81 Wesselink A (1948) Heat conductivity and nature of the lunar surface material. Bull Astron Inst Neth 10:351–363 Wieczorek MA, Huang S (2006) The heat flow of the Moon revisited. EPSC, Berlin, p 448 Wieczorek MA, Phillips RJ (2000) The “Procellarum KREEP Terrrane”: implications for mare volcanism and lunar evolution. J Geophys Res 105(20): 417–420 Williams JP, Paige DA, Greenhagen BT, Sefton-Nash E (2017) The global surface temperatures of the Moon as measured by the Diviner Lunar Radiometer Experiment. Icarus 283:300–325 Winter DF, Saari JM (1969) A particulate thermal model of the lunar soil. Astrophys J 156:1135–1151 Zhiguo M, Yi X, Cai Z, Huang H (2014) Influence of topography on simulated surface temperature. Adv Space Res 54(10):2131–2139

Topographic Studies of the Moon Wayne Bailey Association of Lunar and Planetary Observers, Sewell, NJ, USA

Introduction Scope The narrowest definition of topography is the study of surface position and relief, although it is commonly broadened in a variety of ways. In this entry, in addition to the shape, position, and elevation of lunar surface features, we will include albedo features, since they are visible on images and can assist with interpretation of structural features. Examples will emphasize nearside features, although analogous features are found on the farside. Interpretation of the topography is also limited to what can be inferred from the topography alone.

Topographic Studies of the Moon

Lunar Topographic Structures Lunar topography can be described at several levels. The coarsest level is simply the overall shape of the moon. The next level is the distinction between the rugged highlands and the smoother, and generally lower, mare material. This is mainly a distinction between exposed lunar crust and lava-covered regions. The most obvious topographic features are the maria, which are easily recognized by their dark, smooth appearance. These are large, lava-filled basins, whose smoothness results from their formerly liquid surface. Their dark appearance results from their derivation from subsurface magma which differs from the composition of the crustal rocks. Craters are the iconic lunar features. Their size ranges from the largest mare basins to tiny pits visible only on high-resolution spacecraft images and microscopic pits seen on rocks retrieved from the lunar surface. Most are basically circular, although polygonal approximations to circular are commonly found. Although not common, elongated craters (e.g., Messier) do exist, which are attributed to low-angle impacts. Nearly all craters were formed by hypervelocity impact; only some of the smallest craters may have formed by volcanic processes. Crater topography varies systematically with size. Small craters are simple smooth bowl-shaped depressions with a raised rim (e.g., Linné). Slightly larger craters show a small flat floor (e.g., Chladni). Moderatesized and larger craters are surrounded by a terraced wall and a flat central floor with a central peak or cluster of peaks (e.g., Eratosthenes). Their depth-to-diameter ratio also is smaller than the nearly hemispherical small craters. Even larger craters, often referred to as basins, again lack central peaks (e.g., Archimedes) The largest craters, which include the mare basins, show a multiple ring structure, giving the appearance of multiple, concentric rims. Mare Orientale is the classic example of a multi-ring basin. Since it is only partially flooded, multiple rings are easily visible. In addition, some variations of the basic crater form have been recognized. Ray craters with bright streaks radiating from the crater may be

Topographic Studies of the Moon

the best known (e.g., Copernicus). The rays are formed by fresh, bright material excavated from the parent crater and bright secondary craters created by the ejecta. An inner, dark halo of ejecta often surrounds the crater. A few ray craters exhibit dark rays and a bright halo, such as Dionysius. This results from formation on dark mare thinly covered by lighter highland material. The bright, freshly exposed material gradually darkens on exposure to the space environment, so rays eventually fade from view, making them an indicator of young features. Their rate of fading appears to differ depending on composition, however. Concentric craters exhibit a raised inner ring concentric with their rim. Their origin is not well understood. Dark haloed craters are small craters surrounded by a roughly circular dark halo several times their diameter. Examples are found on the floor of Alphonsus. Those occurring within large craters appear to be pyroclastic deposits surrounding a volcanic vent. Similar appearing dark haloed craters that occur on mare material are impact craters that excavated underlying darker mare material. Floor-fractured craters are large craters whose floor has been uplifted and broken by magma injected from below. Posidonius is one example. Crater chains are simply a series of aligned craters. Some chains are secondary craters formed by ejecta from a larger crater (e.g., Rheita Valley); others may result from binary or fragmented impactors, and some are pits formed by collapse of lava tubes. Most lunar mountain ranges, such as the Apennines and Jura Mountains, are the remnants of basin rims. There are also isolated peaks, usually projecting out of a mare surface (e.g., Mons Pico) and a few isolated ridges, such as the Agricola Mountains. The mound named Mons Rumker appears distinctly different from other features labeled “mountain” and probably results from subsurface magmatic structures. Mountain seems to be a heterogeneous category that simply indicates a positive relief feature. Domes are small, low hills, often with a summit pit. Their slope angle is small so they are only visible under low-angle illumination. They are thought to be volcanic vents. Well-known domes can be found near Milichius and Hortensius. An

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extensive area of volcanic features, including domes, is known as the Marius Hills, adjacent to the crater Marius. Linear features occur in both positive and negative relief forms. Rilles are narrow channels. Rilles that are concentric to the rim of a mare or crater floor are tension faults created by subsidence of the basin center. Examples can be found around the edges of Mare Humorum or Serenitatis. Sinuous rilles were formed by lava flows, either as the surface channel along which lava flowed (e.g., Schröter’s Valley which drained lava from the Cobra Head) or as lava tunnels whose roof subsequently collapsed (such as the Hyginus Rille). Other rilles do not seem to fit either category. For example, a rille in Hesiodus parallels the west wall but then continues north across the north wall to Mare Nubium. There are also rilles such as the Sirsalis Rille that is radial, rather than concentric, to Oceanus Procellarum. Mare (or wrinkle) ridges are positive relief features that are also created by faults, but in this case, compression creates a low-angle fault. Compression causes one side to override the other forming a low ridge. Ridges are common on all maria. Best known is the Serpentine Ridge in Mare Serenitatis. Finally, scarps are faults in which one side is vertically displaced relative to the other side. The Straight Wall is one example, for which the surface expression is actually a fairly gentle slope. The Altai Scarp is another, more dramatic example, which is part of the rim of the Nubium multi-ring basin. Low relief features, such as rilles and ridges, are best seen under low-angle illumination, near the terminator, when shadows increase their contrast. Since solar illumination is oriented east–west, shadows are most effective for features oriented north–south. Albedo features (differences in reflectivity) are not structures, but they are visible, so I consider them topographical objects. In many cases, albedo differences correlate with structural features, as for example the darkness of mare compared to the highlands or dark haloes around craters. Albedo differences can only be isolated at full moon, when unresolved shadow effects are absent. At other phases, the observed intensity results from a convolution of the albedo with a

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phase function which depends on the small-scale surface structure. Differences of albedo with wavelength result in color differences which are related to composition. Lunar colors are faint and differences small, so the moon appears mostly gray. The most colorful region is the area around Aristarchus. Lunar albedo features include shadings on the surface of the maria, which sometimes delineate different lava flows, and Reiner gamma, a bright swirl on Oceanus Procellarum.

Techniques Telescopic Single images provide two-dimension position and brightness information. When combined with knowledge of the illumination direction, elevation information can be extracted, providing a third dimension. Shadow length provides a direct, geometric measurement of local elevation differences. Measurement of the same feature under different solar elevation allows an estimate of the local slope. An indirect method using brightness variations to determine slope angle has been extensively used by the Geological Lunar Research group in Italy to measure dome slopes (Lena et al. 2013). Multiple images taken simultaneously at different wavelengths provide some compositional information through spectral reflectivity differences. This is useful for mapping areas with similar properties. However, diagnostic spectral features of minerals occur mostly in the infrared and usually are indicative of mineral groups, not individual minerals (Serventi 2015). Several spacecraft, most notably Clementine and the Moon Mineralogy Mapper on Chandrayaan-1, used filters specifically designed for identifying the lunar surface composition (Zhang 2015). Simply increasing the color saturation of a properly balanced color image will show lunar colors. Multiple images taken at different lunar phases, and properly calibrated, allow determination of the albedo and phase function. The phase function is determined by the particle size distribution, small-scale slope distribution, and

Topographic Studies of the Moon

shadowing in unresolved crevices (Hapke 1966, 2012). Crater counts have been used extensively to determine exposure ages for the lunar surface. The principle is simply that the longer a surface has been exposed, the more crater forming impacts will have occurred. The age found is thus the time since the most recent resurfacing. Unless the cratering rate is known, only the sequence of formation is found. While simple in principle, there are several factors that complicate the analysis. The cratering rate is not known a priori and varies with both time and crater size. In the oldest, heavily cratered regions, the saturation effect of new craters obliterating older ones limits its usefulness. Dating of samples returned from the moon, whose origin is identifiable, provides absolute age reference points to calibrate the relative age methods. The superposition principle also helps determine sequences of events. A structure or object lying on top of or modifying another must be younger. Thus, the lava that floods a crater must be younger than the crater. Ejecta lying on a mare had to come from an impact that was more recent than the mare flooding event. Alignments can also help identify the source of a structure. Chains of secondary craters point toward the source impact’s location. Linear gouges point toward the basin-forming impact that created the ejecta that produced the gouges. Spacecraft Spacecraft in lunar orbit introduce some additional capabilities (Rennilson 2015), in addition to the telescopic techniques which are also applicable to spacecraft images with higher spatial resolution than ground-based images. One new capability of spacecraft imagery is the ability to determine elevation by stereo-imagery since images taken at different points along the orbit (or on different orbits) view surface points from significantly different directions. Radar or Lidar altimetry allows direct measurement of elevation. This has essentially replaced all other methods for absolute elevation determination. However, the spatial resolution of current

Topographic Studies of the Moon

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digital elevation models is significantly poorer than direct images. Spacecraft have also allowed the precise measurement of the lunar gravitational field which is determined by the mass distribution in surface and subsurface structures. From this, subsurface structures have been identified.

dated. A more recent summary is Hiesinger and Head (2006) which concentrates on new data obtained by spacecraft. Spudis (1993) provides a description of crater formation as an introduction to his main subject, multi-ring basins, the largest craters. Stöffler et al. (Stöffler et al. 2006) describe the use of crater counts to determine lunar ages, along with other age-dating techniques.

Data Sources and Software

Amateur Images Amateur images of the moon are collected and archived by the Association of Lunar and Planetary Observers. Contact information for the appropriate coordinators can be found on the ALPO website at http://alpo-astronomy.org/.

Atlases Two atlases nicely complement each other. Rukl’s Atlas of the Moon (Rukl 2004) is a traditional, hand-drawn atlas, which is nicely complemented by Wood and Collins’ 21st Century Atlas of the Moon (Wood and Collins 2012) which uses labeled photographs. Virtual Moon Atlas, which includes additional capabilities besides being an interactive atlas, is a free program available for download at http://www.ap-i.net/avl. The US Geological Survey hosts the International Astronomical Union’s Working Group for Planetary System Nomenclature’s website Gazetteer of Planetary Nomenclature (lunar page at http:// planetarynames.wr.usgs.gov/Page/MOON/target) which is particularly useful for locating obscure features since it produces images with all approved names. Books Hartmann’s planetary science textbook (Hartmann 2005) provides a good starting point for students or researchers from other fields. Wood’s The Modern Moon (Wood 2003) also provides a very readable introduction to the lunar surface. The most extensive, single published topographic study of the moon is Schultz’s monograph (Schultz 1972) which contains morphological descriptions and interpretations of Lunar Orbiter photographs. Although now somewhat dated, it is still valuable for its detailed examination of numerous features. Mutch (1970) primarily addresses interpretation as compared to Schultz’s descriptive approach. Wilhelms’ Geologic History of the Moon (Wilhelms 1987) includes newer data than Mutch’s volume, but both are now somewhat

Solar System Exploration Research Virtual Institute The Solar System Exploration Research Virtual Institute (SSERVI, http://sservi.nasa.gov/), formerly the NASA Lunar Science Institute, is the most comprehensive source for information about recent lunar projects. It includes links to many projects’ websites. Lunar and Planetary Institute The Lunar and Planetary Institute (LPI, http:// www.lpi.usra.edu/) maintains an online archive of digitized photos and atlases, both ground and space based. Many other resources for researchers and educators are available here, including digitized books and electronic access to publicly available journals. SAO/NASA Astrophysics Data System The Astrophysics Data System (http://adsabs. harvard.edu/), maintained for NASA by the Smithsonian Astrophysical Observatory, provides the ability to search the electronic versions of journals.

Software Lunar Terminator Visualization Tool LTVT is a program that can display and analyze digital elevation models as well as images. It is available for download at http://ltvt.wikispaces.

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com/LTVT. It is capable of adjusting the viewpoint for images and also the illumination direction for DEMs, as well as performing measurements. Ephemeris Data JPL’s Horizons system, available at http://ssd.jpl. nasa.gov/horizons.cgi, provides a very flexible ephemeris generator. For those who need or prefer a local program, the US Naval Observatory’s Multiyear Interactive Computer Almanac (MICA) is available for purchase from http:// www.willbell.com/almanacs/almanac_mica.htm.

References Hapke B (1966) An improved theoretical lunar photometric function. AJ 76:333–339 Hapke B (2012) Theory of reflectance and emittance spectroscopy. Cambridge University Press, New York, pp 303–338 Hartmann WK (2005) Moons & planets. Brooks/ColeThomson Learning, Belmont Hiesinger H, Head JW (2006) New views of lunar geoscience: an introduction and overview. In: Joliff BL, Wieczorek MA, Shearer CK, Neal CR (eds) New views of the Moon, vol 60, Reviews in mineralogy and geochemistry. Mineralogical Society of America, Chantillly Lena R, Wöhler C, Phillips J, Chiocchetta MT (2013) Lunar domes, properties and formation processes. Springer, Milan Mutch TA (1970) Geology of the Moon. Princeton University Press, Princeton Rennilson J (2015) Surveyor Imagery. In: Cudnik B Encyclopedia of Lunar Science. SpringerReference (http:// www.springerreference.com/index/chapterdbid/ 399736) Rukl A (2004) In: Seronik G (ed) Atlas of the Moon, revised updated edition. Sky Publishing, Cambridge Schultz P (1972) Moon morphology. University of Texas Press, Austin Serventi G (2015) Laboratory Analysis (Reflectance Spectra) of Terrestrial Analogues. In: Cudnik B Encylopedia of Lunar Science. SpringerReference (http://www.springerreference.com/index/chapterdbid/ 399733) Spudis PD (1993) The geology of multi-ring impact basins. Cambridge University Press, Cambridge Stöffler D, Ryder G, Ivanov BA, Artemieva NA, Cintala MJ, Grieve RAF (2006) Cratering history and lunar chronology. In: Joliff BL, Wieczorek MA, Shearer CK, Neal CR (eds) New views of the Moon, vol 60, Reviews in mineralogy and geochemistry. Mineralogical Society of America, Chantillly

Total Lunar Occultations Wilhelms DE (1987) Geologic history of the Moon, USGS prof. paper 1348, U.S. Government Printing Office, Washington, DC. Scanned digital version available at http://ser.sese.asu.edu/GHM/ Wood C (2003) The modern Moon: a personal view. Sky Publishing, Cambridge Wood C, Collins M (2012) 21st century atlas of the Moon. Lunar Publishing, UIAI, Wheeling Zhang W (2015) Estimate of Lunar TiO2 and FeO with M3 data. In: Cudnik B Encyclopedia of Lunar Science. SpringerReference (http://www.springerreference. com/index/chapterdbid/399718)

Total Lunar Occultations Richard Nugent International Occultation Timing Association, Dripping Springs, TX, USA

Introduction Occultation observing is a fun venture that almost anyone with minimal equipment and astronomy background can do and one which can lead one toward a potential career in astronomy. This chapter is intended to describe the basics of lunar occultations.

Basics of Occultations An occultation occurs when a solar system body passes in front of a more distant one or a star, partially or totally hiding it, momentarily blocking its light. Each occultation can be seen only at the right time and from a limited part of the Earth. As an example, when an airplane flies directly between an observer on the Earth’s surface and the Sun, its shadow passes right over the observer blocking the Sun’s light momentarily. The airplane has occulted the Sun. Now move the observer just 100 meters away, and the airplane’s shadow would miss the observer entirely. If the observer had arrived 30 s later to that position, he or she would not see the occultation of the Sun by the airplane and have a miss at that location due to being late. Thus occultations are both time and position dependent. The occultation observer

Total Lunar Occultations

must be at the right place at the right time to be in the occulting body’s shadow. Two primary examples of occultation astronomy include lunar occultations and asteroid occultations. Other solar system objects produce other types of occultations, for example, planets, comets and trans-Neptunian Objects (TNOs), Kuiper belt objects (KBOs), and others. These objects can occult stars and one another as well. These latter objects are considered similar to asteroid occultations but are much more difficult to observe.

Scientific Uses of Occultations The news media frequently carries articles and stories about discoveries made with advanced equipment found at major observatories and space probes. In this day of enormous radio telescopes, space missions, the Hubble Space Telescope, the Chandra X-ray telescope, the Space Infrared Telescope Facility (now known as the Lyman Spitzer Observatory), and planetary space probe missions, is an occultation timing with a small telescope of scientific value? The answer is a resounding YES. Occultation timings are one area where interested amateurs can and do make valuable contributions. Occultation observations and an improved lunar ephemeris have been used to determine the location of the celestial equator and the zero-point of right ascension (equinox). These are important corrections to the coordinate system used for stellar reference. Lunar occultation data were given the highest weight for the determination of the fundamental catalogues FK4, FK5, and FK6 stellar reference frame. The analysis of occultation data gathered over a period of many years has allowed astronomers to refine our knowledge of the motion of the Earth, the precession of the North Pole, and the secular motion of the obliquity of the ecliptic (tilt of the Earth’s axis to the equator). Since 1969, lunar orbital parameters have been accurately determined from laser ranging to the retroreflector arrays placed on the Moon by Apollo astronauts and a Soviet spacecraft. Occultation and solar eclipse observations made

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before 1969 are used to study the long-term motion of the Moon. The longer baseline of these observations still gives occultation data an advantage for measuring the Moon’s secular deceleration in ecliptic longitude. This deceleration is caused by an exchange of angular momentum between the rotation of the Earth and the lunar orbit via oceanic tides amounting to 2300 (23 seconds of arc) per century. Recent studies indicate that the tidal deceleration should be about 2800 per century. The difference, an acceleration of 500 per century squared, may be caused by a changing gravitational constant, according to some cosmological theories. Other studies indicate that the difference is smaller and possibly negligible. Occultation data can accurately relate the lunar motion to the stellar reference frame, and the data have proven extremely useful for refining the latter. Lunar occultations have been analyzed to derive average stellar proper motions, and from these the Oort parameters (Oort’s constants) of galactic rotation have been determined. Hundreds of double stars have been discovered by the lunar occultation technique. Occultation observations have been used for hundreds of years by sailors to determine time and their position at sea. Modern occultation observations are routinely used to refine the orbit of the Moon, analyze the positions of stars and the coordinate system they represent, detect new stellar companions, pinpoint the position of X-ray and radio sources, determine the size and shape of lunar mountains, determine stellar diameters, and the recent hot area of determining the size and shape of asteroids in our solar system. In 1985, Pluto’s atmosphere was discovered by the occultation technique. In March 1977, the occultation of a bright star by the planet Uranus resulted in the discovery of its ring system. This ring system might actually have been seen indirectly by the discoverer of Uranus, William Herschel, as he noticed faint stars dim as the planet passed close by. Occultation observations are fun to observe. There is perhaps nothing more exciting than watching a star vanish and return from behind a lunar mountain or to see the star disappear for several seconds as an asteroid passes in front of

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Total Lunar Occultations

Observing lunar occultations tends to be fun but not as important scientifically as it was in the last century. However these events do provide novice observers with a training tool from which they can make timings using visual means or more advanced methods such as with video. Lunar occultations are more easily observed by those with small telescopes. Figure 1 shows lunar occultation geometry. Lunar occultations are ideally suited for small telescopes since they do not require locating a

faint star in the sky. To observe a total lunar occultation of a star, simply find the Moon, and locate the position along the edge of the lunar surface where the star is predicted to be as the Moon moves in front of it. The star should then be easily visible. Since the Moon moves in the easterly direction across the sky, the first 2 weeks following New Moon are the best time to observe a star’s disappearance, since the leading edge of the Moon is dark and not illuminated by the Sun. Similarly, the 2 weeks before New Moon are the best time to observe the reappearance of a star as the Moon uncovers it since the trailing edge of the Moon will not be illuminated by the Sun and will also be dark. Beginners should start by observing total lunar occultations (which occur nightly from nearly everywhere on Earth) and then graduate to observing grazing occultations at the north or south poles of the Moon. Grazing occultation geometry is shown in Fig. 2 and is covered in another chapter in this series.

Total Lunar Occultations, Fig. 1 Lunar occultation. As the Moon moves in its orbit, its projection on Earth results in a total lunar occultation of the star for those

observers between points C and D, while a grazing lunar occultation of the star occurs at points C and D or anywhere along lines N and S

it. Anyone with a small telescope, tape recorder or camcorder, and shortwave radio can make valuable scientific observations to help determine the size and shape of asteroids and to aid in new discoveries about these mysterious objects, including some of the elusive small Moons that orbit them.

Lunar Occultations

Total Lunar Occultations, Fig. 2 Left – lunar grazing occultation geometry. Projection of Moon’s edge on Earth. Right – as the Moon moves, the star disappears/reappears as it passes behind very specific lunar features

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What You Need in the Way of Knowledge and Equipment To observe a total lunar occultation, you will need: (a) Telescope (b) Recorder that can record your voice, either cassette or digital voice recording device (c) Shortwave radio that can pick up time signals at either 5, 10, or 15 Mhz (North America) (d) Predicted time of the occultation event at your location You will begin by observing a star undergoing a lunar occultation, then recording it, and consulting IOTA to reduce and report your data. If you enjoy this observation, you can then advance to the next phase which is video recording of the occultation, but this will require an investment of perhaps USD 300–500 more or less to obtain a video camera, battery, digital video recorder (usually from eBay), associated cables, adapters, and a GPS video time inserter. The GPS-based video time inserter will identify and record the observing site location’s latitude

and longitude. The GPS time (Universal Time, UT) will be overlaid on the video in real time to 1/100 of a second or better.

What an Observer Should Expect to See A total lunar occultation will result in the star disappearing behind the lunar limb. Tens of minutes later the star will reappear from behind the Moon. Step events (where the star’s brightness does not drop to zero) may also be detected. Such events may be the sign of a double star and/or an unknown double star. When the Moon occults a double star, there will a step decrease in brightness as shown in the light curve shown in Fig. 3.

The Need for Accurate Timing Without having a record of accurate time, your observations are relatively useless for scientific contribution. If you are calling out your times into a voice recorder while watching through the telescope, your reaction time (difference between the moment you saw the event and when you called it out) varies from a fraction of a second

Analyzed film name [Mini-DVR-640x510.avs] Photometry in each Frame

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to several seconds depending upon the level of experience of the observer and other factors. Occultation timings are generally needed to millisecond accuracy (0.033 s) or better. This is best accomplished using video recordings with a GPS time inserter. Beginning visual observers can use a tape recorder and voice “call outs” to record an observation, but one will soon easily see upon attempting to reduce data from several closely spaced observers that there are significant errors in the visual observing/timing process. The best method of recording an occultation of any type is with a sensitive video camera, a GPS time inserter, and a digital or analog camcorder – not a computer. While a laptop computer seems like an ideal method to record video/timing with the multitude of software applications/video capture cards, there are inherent problems and processing delays with these setups that can lead to timing errors of 0.10 s and larger.

Total Lunar Occultations

IOTA The International Occultation and Timing Association (IOTA) is a volunteer organization which predicts, gathers, analyzes, and publishes observations of occultations. It has many resources available to help you get started in this fascinating adventure.

Cross-References ▶ Kaguya (SELENE) Mission ▶ Lunar Reconnaissance Orbiter (LRO) Mission ▶ Topographic Studies of the Moon

References David Dunham (1975) IOTA. http://www.occultations.org

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Ultraviolet Radiation Caitlin Ahrens NASA Goddard Space Flight Center, Greenbelt, MD, USA

Definition The lunar surface has a complex plasma environment due to incoming solar wind flux and charged micron-sized dust particles and the presence of solar ultraviolet (UV) radiation. UV resonance fluorescence of solar radiation provides a sensitive detection of major elemental constituents, such as H, H2, O, C, N, and noble gases in the tenuous daytime lunar atmosphere. Photoemission due to the solar UV radiation is the dominating process in the charging environment, creating a photoelectron “sheath.” Such detections were observed by LADEE UV spectrometer and the Apollo 17 UV spectrometer.

Lunar Dusty Plasma On the sunlit side of the lunar surface in the solar wind, photoelectric charging due to UV radiation and the consistent collection of solar ions and electrons dominate the lunar plasma environment (Poppe and Horanyi 2010; Örger et al. 2017). This sunlit side is expected to charge positively. Since © Springer Nature Switzerland AG 2023 B. Cudnik (ed.), Encyclopedia of Lunar Science, https://doi.org/10.1007/978-3-319-14541-9

the photoelectron current dictates the charging process of the sunlit lunar side, it is interesting to note that the photoemission properties of the lunar soil are diverse (Lisin et al. 2015). The sources can also be used to determine the surface potential and electric field, which can also vary as the current sources alter with time. Örger et al. (2017) suggest that the lunar horizon glow is produced by scattering of the sunlight by electrically charged dust grains, though the physical mechanism of this process is still under investigation. UV charging of the grains may be caused by several processes (Abbas et al. 2007), which include the following: (i) ion or electron collisions on the nightside, leading to negatively charged grains with F > Cl.

Abundance of Volatiles in urKREEP Regrettably, our lunar sample collection does not contain a sample directly related to urKREEP, similarly to the way basaltic volcanic products relate to the lunar mantle via partial melting. The composition of urKREEP can only be estimated from specific lunar samples with KREEP-rich

geochemical signatures. The highlands Mg-suite rocks are KREEP-rich igneous cumulates whose parental magmas involved interaction of earlyformed Mg-rich LMO cumulates with the anorthosite crust and urKREEP liquids (e.g., Elardo et al. 2011), the latter dominating the volatile signature of Mg-suite melts since early-formed magnesian LMO cumulates should have contained very little volatiles. The volatile abundances measured in apatite in highlands Mg-suite rocks suggest that the melts from which they crystallized generally contained more Cl than H2O and F (Fig. 2), indicating that the dominant volatile in urKREEP was Cl. A couple of studies have tried to constrain absolute volatile abundances in urKREEP. Based on Apollo and Luna soil bulk analyses, Treiman et al. (2014) proposed that urKREEP contained ~660 ppm F and ~150 ppm Cl, and Hui et al. (2013) suggested it contained up to 1.4 wt% H2O from the analysis of plagioclase in an Apollo 16 ferroan anorthosite. Combining these estimates yields H2O >> F > Cl, which is not consistent with the relative volatile abundances determined from apatite (Cl > H2O ~ F). As argued by McCubbin et al. (2015b), the F abundance estimate for urKREEP is likely the most robust since (i) it is least susceptible to volatilization compared to H2O and Cl, (ii) only the nonleachable Cl fraction was used to estimate the Cl urKREEP

Water in the LMO

content by Treiman et al. (2014), and (iii) partitioning of water in anorthitic plagioclase at lunar conditions is not well constrained. Combining the F content estimated for urKREEP with volatile abundances of highlands apatite and apatite/melt partitioning data, McCubbin et al. (2015b) have proposed that urKREEP could have contained around 660 ppm F, 1100–1350 ppm Cl, and 300–1250 ppm H2O.

Abundance of Volatiles in the Bulk Silicate Moon Based on this estimate of volatile abundances in urKREEP, it is possible to calculate the volatile abundances in the bulk silicate Moon (BSM), which correspond in fact to those in the fully molten LMO. The Cl depletion of the lunar mantle, and its reciprocal enrichment in urKREEP compared to H2O and F, is consistent with volatile abundances having been controlled by the partitioning behavior of H2O, F, and Cl between LMO melts and early-crystallized olivine and pyroxene. Therefore, it is possible to estimate volatile abundances of the LMO before it started to crystallize using appropriate LMO crystallization models. Using their urKREEP estimate (300–1250 ppm H2O, 660 ppm F and 1100–1350 ppm Cl) and assuming that 0.5 % of residual liquid remained trapped in the mantle cumulate pile, McCubbin et al. (2015b) proposed that the BSM contains around 3–13 ppm H2O, 7 ppm F, and 11–14 ppm Cl. Based on the volatile abundances measured in picritic glasses, Hauri et al. (2015) proposed that the BSM contains 133–292 ppm H2O, 4.5–5.4 ppm F, 0.14–0.20 ppm Cl, and 79 ppm S. The H2O and Cl abundances estimated by Hauri et al. (2015) are, respectively, higher and lower than those proposed by McCubbin et al. (2015b), while the F content of the BSM is similar in both estimates. However, McCubbin et al. (2015b) postulated that no F had been mobilized from lunar soils over the geological history of the Moon, an assumption that might not be entirely true. As a result, they proposed an upper bound on the volatile abundance of the BSM by

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considering a chondritic F abundance of 60 ppm for the BSM. In this scenario, urKREEP liquids would have contained 0.26–1.09 wt% H2O, 0.55 wt% F, and 0.98–1.20 wt% Cl, resulting in BSM volatile abundances of 27–114 ppm H2O, 60 ppm F, and 100–123 ppm Cl (McCubbin et al. 2015b). This improvement between the match of BSM H2O contents between the two estimates, however, increases the discrepancy regarding the BSM Cl content. The inconsistency between estimates of the volatile abundance of the lunar interior based on different types of samples may either be indicative of a heterogeneous distribution of volatiles among different indigenous reservoirs (e.g., Robinson and Taylor 2014) or be symptomatic of our fragmented understanding of the origin and petrogenesis of the lunar lithologies investigated. For example, in McCubbin’s upper bound scenario, the cumulate mantle contains 1.2–5.3 ppm H2O, 4.5 ppm F, and 2.3–2.9 ppm Cl. Such water contents are at the lower end of estimates based on volatile abundances measured in apatite in mare basalts and in picritic glasses. This could indicate that either the H2O content estimates based on mare basalts and picritic glasses are robust but represent individual source regions in a heterogeneous mantle, or that many estimates of the H2O content of the lunar mantle have been biased towards the wettest samples. The latter is not unrealistic considering that partial melting in the lunar interior likely preferentially affected the wettest source regions since addition of H2O reduces solidus temperatures in silicate systems.

Source(s) and Timing of Delivery of Indigenous Lunar Water The isotopic composition of the hydrogen in water, expressed by the D/H ratio (or by the δD notation, which corresponds to the deviation in permil compared to the average D/H ratio of the Earth’s ocean), provides important clues regarding its origin(s). Indeed, D/H ratios of different Solar System objects vary by about one order of magnitude, ranging from ~25  10 6 for the proto-sun to ~200–400  10 6 for Oort Cloud

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Water in the LMO

Water in the LMO, Fig. 4 D/H ratios of highlands and mare basalt apatite compared to those of different Solar System objects

cometary ices, with water in the Earth and in most carbonaceous chondrite objects clustering around 150  10 6 (e.g., Robert 2006) (Fig. 4). In situ analysis of apatite and picritic glasses has provided remarkable information regarding the H isotope composition of lunar indigenous water, but the interpretation of these data is still a topic of great debate. Greenwood et al. (2011) published the first D/H measurements in lunar apatite and showed that D/H ratios in the large majority of mare basalt apatite are much higher than those measured in terrestrial rocks, with δD values up to ~1000 ‰, which led these authors to suggest that the lunar interior contains a D-rich reservoir originating from the delivery of cometary H2O (Fig. 4). Subsequent studies confirmed the D-rich nature of water in apatite in most mare basalts (Barnes et al. 2013; Tartèse et al. 2013), but proposed a different interpretation in which water in mare basalts was initially characterized by D/H ratios similar to those measured in terrestrial rocks and in the majority of carbonaceous chondrites (δD ≈ 0  200 ‰), the elevated D/H ratios resulting from the preferential degassing of the lighter H2 molecule compared to HD during crystallization of the lavas on the lunar surface (Tartèse and Anand 2013; Tartèse et al. 2013). Interpretations of D/H lunar signatures are further complicated by possible contamination by solarwind H (δD ≈ 1000 ‰) implanted in lunar soils that could be assimilated by hot lunar magmas and

by the effect of spallation reactions due to interaction with galactic cosmic-rays, producing D in materials residing close to the lunar surface. Effects of spallation on D/H ratios are particularly noticeable in picritic glasses, which contain much less water than apatite in most mare basalts. Once corrected for spallation effects, δD values of water in high-Ti picritic glasses appear to cluster in the range ~0  200 ‰ (Saal et al. 2013; Füri et al. 2014). In these high-Ti picritic glasses, Saal et al. (2013) measured δD values of ~200 ‰ for water trapped in olivine-hosted melt inclusions and argued that water in the mantle source region of these glasses was characterized by D/H ratios similar to those of CI carbonaceous chondrites, just slightly heavier than terrestrial D/H ratios. Finally, analysis of apatite in three highlands samples yielded D/H ratios much lower than for mare basalts (Fig. 4), with average δD values ranging between about 200 ‰ and 0 ‰ (Barnes et al. 2014), which these authors interpreted as representative of the δD signature of urKREEP (and hence of the BSM). This is consistent with the average δD value of 130  50 ‰ measured in apatite in two KREEP-rich basalts (Apollo sample 72275 and lunar meteorite Northwest Africa 773) by Tartèse et al. (2014). Most of the studies carried out on lunar apatite and picritic glasses since the pioneering work of Greenwood et al. (2011)

Water in the LMO

seems to converge towards a consensus view that indigenous lunar water generally has a similar H isotope composition to that of most types of carbonaceous chondrites and of the Earth, even though the presence of water reservoirs with heterogeneous D/H signatures cannot be excluded (Robinson and Taylor 2014). The delivery of lunar indigenous water by asteroidal objects similar to most types of carbonaceous chondrites is also consistent with the nitrogen isotope characteristics of lunar mare basalts (Mortimer et al. 2015; Füri et al. 2015). The detection of significant amounts of water (and other volatiles such as F and Cl) in early products of lunar differentiation such as ferroan anorthosites and Mg-suite norites (Hui et al. 2013; Barnes et al. 2014) suggests that the LMO contained water at the time of, or shortly after, its formation. As argued in detail by Hauri et al. (2015), the formation of the Moon in the context of the Giant Impact hypothesis can be decomposed in three major stages characterized by very different timescales, offering time windows during which water (and other volatiles) can be added to, or lost from, the Moon. Ejection of material following the Giant Impact and evolution of the proto-lunar disk of magma and vapor both offer possibilities for loss of water and other volatile species on timescales ranging from days to hundreds of years. It has been proposed that hydrodynamic escape of light volatile species such as H2 and H2O could occur during this proto-lunar disk phase (Pahlevan and Stevenson 2007), although recent modeling studies have suggested that this phenomenon might be in fact limited as long as the vapor phase is dominated by silicate vapor (Visscher and Fegley 2013; Nakajima and Stevenson 2014). On the other hand, coalescence of materials that will eventually form the Moon outside the Roche radius from a thin magma disk likely incorporated very little amounts of the vaporized materials, containing the vast majority of the volatiles, since most of these materials would have been gravitationally bound to the proto-Earth (Hauri et al. 2015). Therefore, the Moon likely formed with very little water (as well as other volatiles). Cooling and crystallization of the LMO offers a time window

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of 10–200 Ma before crystallization of a rigid lithosphere for the delivery of a mass of volatilerich, carbonaceous chondrite-like, material to the Moon, which would be consistent with fluxes estimated for the tail-end of planetary accretion (O’Brien et al. 2014). In addition to this “hot start,” Hauri et al. (2015) also argued that a “cold start” scenario, in which the Moon accreted from material in the proto-lunar disk that remained largely solid from which they inherited most of their volatiles, was a viable possibility.

Future Directions As discussed above, estimates of the abundances of water (and of other volatiles such as F and Cl) in the BSM, and hence in the LMO, show slight differences depending on the different types of samples investigated. This highlights the importance of using sets of lithologies as diverse as possible in order to further constrain the volatile inventory of the BSM, but also of investigating in detail the petrogenetic history of the studied samples in order to identify processes that could have resulted in gains or losses of volatiles. Analysis of volatile contents of nominally anhydrous minerals such as pyroxene and plagioclase could provide additional, fundamental constraints on the volatile inventory of the silicate melts from which they crystallized since partitioning behavior of volatiles in these minerals is less complicated than in minerals such as apatite. Accurate analysis of volatile contents in nominally anhydrous minerals, and especially of water, will require (i) further developments of SIMS analytical techniques, in order to continue to reduce instrumental background volatile contents, and (ii) the characterization of appropriate and widely available reference materials. Also, future research should focus its efforts on devising experiments relevant for lunar conditions (P, T, fO2) as well as aiming to better quantify the partitioning behavior of volatiles between silicate liquids and both nominally anhydrous minerals and volatile-bearing minerals such as apatite. Regarding isotopic studies aimed at constraining the origin of water in the lunar

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interior, it remains ambiguous whether large H isotope variations observed in some rock types, or even within individual minerals in a rock sample, resulted from magmatic fractionation processes or mixing of different water reservoirs with contrasted D/H signatures. Again, detailed petrological analysis of the studied samples could help place constraints regarding fractionation versus mixing processes. Such constraints could also be provided by combining different isotopic systems (C, N, Cl, or S) on individual samples. Finally, as lunar geologists we should never forget that Apollo samples are invaluable but represent only a limited, and possibly biased, sampling of the Moon. Lunar meteorites, which were launched from random sites, have provided us with an additional and wider sampling of the lunar surface. However, they lack a precise geological context, which can render interpretation of data equivocal. Therefore, additional sampling of the Moon, both laterally and vertically, targeted on areas whose geological interest has been assessed using data acquired by orbiting spacecraft, is highly desirable in order to better understand processes by which the Moon formed and its subsequent geological evolution.

References Anand M, Tartèse R, Barnes JJ (2014) Understanding the origin and evolution of water in the Moon through lunar sample studies. Philos Trans R Soc A 372:20130254 Aubaud C, Hauri EH, Hirschmann MM (2004) Hydrogen partition coefficients between nominally anhydrous minerals and basaltic melts. Geophys Res Lett 31: L20611 Barnes JJ, Franchi IA, Anand M, Tartèse R, Starkey NA, Koike M, Sano Y, Russell SS (2013) Accurate and precise measurements of the D/H ratio and hydroxyl content in lunar apatites using NanoSIMS. Chem Geol 337–338:48–55 Barnes JJ, Tartèse R, Anand M, McCubbin FM, Franchi IA, Starkey NA, Russell SS (2014) The origin of water in the primitive Moon as revealed by the lunar highlands samples. Earth Planet Sci Lett 390:244–252 Beyer C, Klemme S, Wiedenbeck M, Stracke A, Vollmer C (2012) Fluorine in nominally fluorine-free mantle minerals: experimental partitioning of F between olivine, orthopyroxene and silicate melts with implications

Water in the LMO for magmatic processes. Earth Planet Sci Lett 337–338:1–9 Boyce JW, Liu Y, Rossman GR, Guan Y, Eiler JM, Stolper EM, Taylor LA (2010) Lunar apatite with terrestrial volatile abundances. Nature 466:466–469 Boyce JW, Tomlinson SM, McCubbin FM, Greenwood JP, Treiman AH (2014) The lunar apatite paradox. Science 344:400–402 Dalou C, Koga KT, Le Voyer M, Shimizu N (2014) Contrasting partition behavior of F and Cl during hydrous mantle melting: implications for Cl/F signature in arc magmas. Prog Earth Planet Sci 1:26 Elardo SM, Draper DS, Shearer CK (2011) Lunar Magma Ocean crystallization revisited: bulk composition, early cumulate mineralogy, and the source regions of the highlands Mg-suite. Geochim Cosmochim Acta 75:3024–3045 Fogel RA, Rutherford MJ (1995) Magmatic volatiles in primitive lunar glasses: I. FTIR and EPMA analyses of Apollo 15 green and yellow glasses and revision of the volatile-assisted fire-fountain theory. Geochim Cosmochim Acta 59:201–215 Füri E, Deloule E, Gurenko A, Marty B (2014) New evidence for chondritic lunar water from combined D/H and noble gas analyses of single Apollo 17 volcanic glasses. Icarus 229:109–120 Füri E, Barry PH, Taylor LA, Marty B (2015) Indigenous nitrogen in the Moon: constraints from coupled nitrogen–noble gas analyses of mare basalts. Earth Planet Sci Lett 431:195–205 Greenwood JP, Itoh S, Sakamoto N, Warren PH, Taylor LA, Yurimoto H (2011) Hydrogen isotope ratios in lunar rocks indicate delivery of cometary water to the Moon. Nat Geosci 4:79–82 Hauri E, Gaetani G, Green T (2006) Partitioning of water during melting of the Earth’s upper mantle at H2Oundersaturated conditions. Earth Planet Sci Lett 248:715–734 Hauri EH, Weinreich T, Saal AE, Rutherford MC, Van Orman JA (2011) High pre-eruptive water contents preserved in lunar melt inclusions. Science 333:213–215 Hauri EH, Saal AE, Rutherford MJ, Van Orman JA (2015) Water in the Moon’s interior: truth and consequences. Earth Planet Sci Lett 409:252–264 Heiken GH, Vaniman DT, French BM (1991) Lunar sourcebook. Cambridge University Press, New York, 756 pp Hui H, Peslier AH, Zhang Y, Neal CR (2013) Water in lunar anorthosites and evidence for a wet early Moon. Nat Geosci 6:177–180 Jolliff BL, Wieczorek MA, Shearer CK, Neal CR (2006) New views of the Moon, vol 60, Reviews in Mineralogy & Geochemistry. Mineralogical Society of America, Chantilly, 721 pp McCubbin FM, Nekvasil H, Lindsley DH (2007) Is there evidence for water in lunar magmatic minerals? A crystal chemical investigation. Paper presented at

Water in the LMO the 38th Lunar and Planetary Science Conference, Houston, 12–16 Mar 2007, Abs #1354 McCubbin FM, Steele A, Hauri EH, Nekvasil H, Yamashita S, Hemley RJ (2010a) Nominally hydrous magmatism on the Moon. Proc Natl Acad Sci U S A 107:11223–11228 McCubbin FM, Steele A, Nekvasil H, Schnieders A, Rose T, Fries M, Carpenter PK, Jolliff BL (2010b) Detection of structurally bound hydroxyl in fluorapatite from Apollo mare basalt 15058,128 using TOF-SIMS. Am Mineral 95:1141–1150 McCubbin FM, Jolliff BJ, Nekvasil H, Carpenter PK, Zeigler RA, Steele A, Elardo SM, Lindsley DH (2011) Fluorine and chlorine abundances in lunar apatite: implications for heterogeneous distributions of magmatic volatiles in the lunar interior. Geochim Cosmochim Acta 75:5073–5093 McCubbin FM, Vander Kaaden KE, Tartèse R, Boyce JW, Mikhail S, Whitson ES, Anand M, Franchi IA, Wang J, Hauri EH (2015a) Experimental investigation of F, Cl, and OH partitioning between apatite and Fe-rich basaltic melt at 1.5 GPa and 950–1000  C. Am Mineral 100:1790–1802 McCubbin FM, Vander Kaaden KE, Tartèse R, Klima RL, Liu Y, Mortimer JI, Barnes JJ, Shearer CK, Treiman AH, Lawrence DJ, Elardo SM, Hurley DM, Boyce JW, Anand M (2015b) Volatiles (H, C, N, F, S, Cl) in the lunar mantle, crust, and regolith: distribution, processes, sources, and significance. Am Mineral 100:1668–1707 Mortimer J, Verchovsky AB, Anand M, Gilmour I, Pillinger CT (2015) Simultaneous analysis of abundance and isotopic composition of nitrogen, carbon, and noble gases in lunar basalts: insights into interior and surface processes on the Moon. Icarus 255:3–17 Nakajima M, Stevenson DJ (2014) Investigation of the initial state of the Moon-forming disk: bridging SPH simulations and hydrostatic models. Icarus 233:259–267 O’Brien DP, Walsh KJ, Morbidelli A, Raymond SN, Mandell AM (2014) Water delivery and giant impacts in the “Grand Tack” scenario. Icarus 239:74–84 O’Leary JA, Gaetani GA, Hauri EH (2010) The effect of tetrahedral Al3+ on the partitioning of water between clinopyroxene and silicate melt. Earth Planet Sci Lett 297:111–120 Pahlevan K, Stevenson DJ (2007) Equilibration in the aftermath of the lunar-forming giant impact. Earth Planet Sci Lett 262:438–449 Robert F (2006) Solar System deuterium/hydrogen ratio. In: Lauretta DS, McSween HY Jr (eds) Meteorites and the early Solar System II. University of Arizona Press, Tucson, pp 341–351 Robinson KL, Taylor GJ (2014) Heterogeneous distribution of water in the Moon. Nat Geosci 7:401–408 Saal AE, Hauri EH, Rutherford MJ, Cooper RF (2007) The volatile contents (CO2, H2O, F, S, Cl) of the lunar picritic glasses. Paper presented at the 38th Lunar and

1265 Planetary Science Conference, Houston, 12–16 Mar 2007, Abs #2148 Saal AE, Hauri EH, Lo Cascio M, Van Orman JA, Rutherford MC, Cooper RF (2008) Volatile content of lunar volcanic glasses and the presence of water in the Moon’s interior. Nature 454:192–196 Saal AE, Hauri EH, Van Orman JA, Rutherford MJ (2013) Hydrogen isotopes in lunar volcanic glasses and melt inclusions reveal a carbonaceous chondrite heritage. Science 340:1317–1320 Sclar CB, Bauer JF (1975) On the halogen deficiency of lunar apatite. Meteoritics 10:484–485 Smith JV, Anderson AT, Newton RC, Olsen EJ, Wyllie PJ, Crewe AV, Isaacson MS, Johnson D (1970) Petrologic history of the moon inferred from petrography, mineralogy and petrogenesis of Apollo 11 rocks. Paper presented at the Apollo 11 Lunar Science Conference, Houston, 5–8 Jan 1970, pp 897–925 Snyder GA, Taylor LA, Neal CR (1992) A chemical model for generating the sources of mare basalts – combined equilibrium and fractional crystallization of the lunar magmasphere. Geochim Cosmochim Acta 56:3809–3823 Tartèse R, Anand M (2013) Late delivery of chondritic hydrogen into the lunar mantle: insights from mare basalts. Earth Planet Sci Lett 361:480–486 Tartèse R, Anand M, Barnes JJ, Starkey NA, Franchi IA, Sano Y (2013) The abundance, distribution, and isotopic composition of hydrogen in the Moon as revealed by basaltic lunar samples: implications for the volatile inventory on the Moon. Geochim Cosmochim Acta 122:58–74 Tartèse R, Anand M, McCubbin FM, Elardo SM, Shearer CK, Franchi IA (2014) Apatites in lunar KREEP basalts: the missing link to understanding the H isotope systematics of the Moon. Geology 42:363–366 Tenner TJ, Hirschmann MM, Withers AC, Hervig RL (2009) Hydrogen partitioning between nominally anhydrous upper mantle minerals and melt between 3 and 5 GPa and applications to hydrous peridotite partial melting. Chem Geol 262:42–56 Treiman AH, Boyce JW, Gross J, Guan Y, Eiler JM, Stolper EM (2014) Phosphate-halogen metasomatism of lunar granulite 79215: impact-induced fractionation of volatiles and incompatible elements. Am Mineral 99:1860–1870 Visscher C, Fegley B Jr (2013) Chemistry of impactgenerated silicate melt-vapor debris disks. Astrophys J Lett 767:L12 Warren PH, Wasson JT (1979) The origin of KREEP. Rev Geophys Space Phys 17:73–88 Wood JA, Dickey JS Jr, Marvin UB, Powell BN (1970) Lunar anorthosites and a geophysical model of the moon. Paper presented at the Apollo 11 Lunar Science Conference, Houston, 5–8 Jan 1970, pp 965–988

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Weathering Kaushik Mitra Department of Geosciences, Stony Brook University, Stony Brook, NY, USA

Definition/Description Weathering is the physical or chemical breakdown of rocks and minerals occurring in situ due to the effects of water, wind, glacier, and biological agents. The major/primary natural weathering forces on Earth are water and wind. However, weathering on Moon (and other airless bodies) is different than on Earth due to the absence of both these weathering agents. The agents/forces of erosion originate external to the Moon, producing an altogether different process called space weathering. The gradual alteration of materials when they are exposed to the variety of natural processes on the surface of airless bodies like Moon, Mercury, and the asteroids is termed as space weathering. It is the combined effect of all the physical, chemical, structural, and optical changes occurring on materials exposed on the surface of planetary bodies like the Moon that are not protected by an atmosphere or a magnetic Weathering, Fig. 1 The full photograph of the Moon showing areas containing craters with “lunar rays.” (Image adapted from Hawke et al. 2004)

Weathering

field, which makes them vulnerable to the space environment. The interaction of the surface material with the space environment, which includes micrometeoroid bombardment, highly energetic particles (including protons, electrons, and alpha particles), and radiation (gamma rays, X-rays, and ultraviolet light), gives rise to space weathering. The multiple processes acting simultaneously on the surface of Moon due to external forces leads to the in situ modification of lunar surface materials and is collectively termed space weathering on the Moon.

Space Weathering: A Historical Perspective The effects of space weathering were first observed in 1955 in which certain younger craters (like Tycho and Olbers A) (Fig. 1) were associated with ephemeral high-albedo, bright features called “rays (Gold 1955).” Based on his observation, Gold suggested that the surface of Moon must be changing color (and likely other optical properties) over time. We now understand that the disappearance of rays (specifically “immature” rays) due to darkening of the regolith (lunar soil) over time is a direct impact of space weathering (Lucey et al. 2006; Pieters and Noble 2016).

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Weathering, Fig. 2 The various space weathering processes that occur on the surface of airless bodies including the Moon. (Image adapted from: Noble 2004; Pieters and Noble 2016)

This hypothesis was confirmed when samples from the surface of Moon were brought back to Earth by the Apollo astronauts (Fig. 2). It was observed that the optical properties of the lunar regolith differ substantially from fresh pulverized lunar rocks from which the regolith was derived. The natural lunar soils are darker than the freshly pulverized rocks and also exhibit weaker spectral absorption bands that are characteristic of the minerals present in them. The change in optical properties is a direct result of the effect of space weathering on the Moon.

Space Weathering Processes: The Different Agents of Space Weathering Space weathering operates at subgranular and grain-size scale in the lunar soil, which has itself been produced by larger-scale processes (Fig. 3). The agents of space weathering can be broadly classified into two categories, micrometeoroids and radiation.

Impact from Micrometeoroids (Small “Dust” Particles or Debris) The impact of micrometeoroid bombardment on the lunar soil produces the following effects: (a) Comminution or disaggregation: The breaking down of larger particles into smaller particles due to the energy released by impacts. (b) Melting and vaporization: The energy of the impact can melt and vaporize the impactor material as well as the host and can lead to a wide range of possible alteration products. The impact-vaporized material can also be condensed back and mixed with the lunar soil. (c) Agglutinate formation: Agglutinates are aggregates of lunar soil particles (mineral grains, glasses, and even older agglutinates) bonded together by glass. The small amount of heat produced due to micrometeoroid impact on the lunar soil causes small-scale melting and leads to agglutinate formation by the fusion of adjacent grains and the inclusion of minor amounts of iron in them. (d) Contamination: A portion of the impactor material can be retained by the host rock and

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Weathering, Fig. 3 Visible/near-infrared (VNIR) spectra of coarsely particulate lunar basalt and lunar soil collected in the laboratory showing substantial differences in their optical properties as a result of space weathering. The

image illustrates the stark differences between fresh basalt and the mature soil derived from the same rock, demonstrating the role of space weathering on the Moon. (Image adapted from: Pieters and Noble 2016)

can lead to a contamination producing a mixed product that depends upon various parameters (like impactor velocity, impactor and target composition, and impact geometry). (e) Soil gardening: The mixing of previously weathered and processed regolith by repeated impacts of various sizes. (f) Production of nanophase iron metal (npFe0): The micrometeoroid impact can lead to production of very small (4–33 nanometers in diameters) metal iron grains by the reduction of Fe2+. Recent research has demonstrated that space weathering generates nanophase Fe particles with a range of oxidation states (Thompson et al. 2016).

(a) Sputtering: The release and/or rearrangement of surface atoms due to the interaction of surface material with energetic particles like cosmic rays and solar wind ions. The sputtered material can also be condensed back and lead to mixing and contamination. (b) Solar wind implantation: Implantation of ions from the solar wind, solar energetic particles, and galactic cosmic rays. (c) Vitrification: The intense solar UV radiation and the energetic cosmic galactic rays can damage the original crystalline structure of the uppermost surface of grains in the soil. (d) Diurnal thermal cycling: Solar insolation variations can lead to structural fatigue of surface materials and can also alter their composition by radiant heating and sublimation.

Radiation Damage from Electromagnetic Radiation and Energetic Particles The bombardment of energetic particles and radiation can lead to the following processes:

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The Primary Agents of Space Weathering Critical aspects of micrometeoroid impacts are the ways in which they differ from larger impacts. Owing to the very short timescales in which micrometeoroid impacts occur, the processes, results, and implications of micrometeoroid impact are fundamentally different from larger impact processes. The micrometeoroids such as interplanetary dust particles (IDP) are typically less than a millimeter in diameter and they strongly affect the lunar regolith by continuously bombarding the lunar surface. As a result of the impact, the micrometeoroid vaporizes and causes the lunar regolith to both vaporize and melt. A combined rate of erosion due to micrometeoroid impact and sputtering reactions (next section) have been estimated to be ~1–2 millimeters per million years. The energetic particle and X-ray interactions with the lunar surface have a twofold significance for the Moon: (i) physical and chemical alteration of the lunar regolith and (ii) enabling the use of gamma ray, X-ray, and neutron spectroscopic remote sensing of the Moon. The energetic particles in the lunar environment are mainly (i) the solar wind, (ii) solar energetic particles, and (iii) galactic cosmic rays. The interaction of lunar materials with the energetic particles can either lead to their implantation in the lunar materials or can induce nuclear reactions. The relative energies of these energetic particles are lowest in solar wind, intermediate in solar cosmic rays, and highest in galactic cosmic rays. Solar Wind Near the surface of the Moon, the solar wind (a plasma emanating from the Sun) travels at velocities of ~300–800 km/s and have energies of ~1 keV/nucleon. The particles (equal parts electrons and ions) in the solar wind can penetrate up to a few tens of nanometers into the lunar materials and can either remain implanted indefinitely or could diffuse back out into space. Gases sourced from the solar wind (like helium, nitrogen, and argon) have been shown to be implanted in lunar soils and their concentration increases as the regolith becomes more mature. Many noble gas

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isotopes have been used to track the surface exposure time and thus location history of lunar materials. Solar Cosmic Rays Solar cosmic rays are orders of magnitude higher in energy (typically ~1–100 MeV/nucleon and occasionally GeV/nucleon) than the solar wind. They are about 98% protons and can penetrate in the lunar material up to ~1 centimeter into lunar regolith grains. The concentration of cosmogenic nuclides with depth on the Moon have been used to determine the average fluxes of solar protons during the past ~10 million years. Galactic Cosmic Rays The galactic cosmic rays contain particles that are highest in energy, typically in the range of ~0.1–10 GeV/nucleon, and comprise of about 78% protons and 12% alpha particles (doubly ionized helium atoms). These particles can penetrate meters into the lunar surface and produce a slew of numerous secondary energetic particles as a result of their interaction with the lunar materials. The energetic particles are slowed down as a result of the interaction with the lunar materials and can induce radiation damage observed as tracks or thermoluminescence as a result. The higher-energy particles in the solar cosmic rays and the galactic cosmic rays can also induce nuclear reactions to produce a variety of stable and radioactive nuclides like 21Ne and 26Al.

Maturation and Evolution of the Lunar Regolith The lunar regolith is the primary component on the surface of Moon that accumulates and records the principal effects of space weathering. Therefore, it is critical to understand the evolution of the lunar regolith. The degree to which the lunar regolith accumulates space weathering products over time is measured by maturity. It is a relative term and is usually used to describe surface soil on the Moon. Another maturity parameter used is exposure age, which is an estimate of the time duration a rock or

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a soil grain has been exposed to space weathering based on the precise measurements of space weathering products such as solar wind noble gases or cosmic ray tracks. The major impact of space weathering on the physical properties of the lunar regolith involves a combination of three processes: comminution, agglutination, and mixing. Comminution dominates over agglutination at the start when a fresh lunar rock or an impact ejecta block is bombarded with micrometeoroids. This leads to a reduction in the average grain size of the particles with time. As the soil matures, micrometeoroid impacts can utilize the heat generated during the impacts and coalesce the fine-grained materials into coarsegrained, glassy materials called agglutinates. Therefore, lunar regolith develops a higher percentage of agglutinates and finer grained particles with time and becomes more “mature.” Other signs of maturity include a lower albedo, higher concentrations of solar-wind-implanted ions and nanophase Fe particles, and a higher density of energetic particle tracks.

The Effects of Space Weathering The rate and extent of space weathering on any planetary body is a function of primarily three parameters: (i) First, the location of the planetary body in the solar system and the presence or absence of magnetic field around the body. This determines the impact speed and flux of the space weathering agents, the radiation environment, and the temperature. (ii) Second, the type of planetary body, i.e., the composition, size, and texture of the host. (iii) Finally, the duration of space weathering. Based on these three parameters, space weathering can produce varying effects on the optical properties of materials in the lunar surface which can broadly be described as: 1. Overall reduction of reflectance, i.e., the albedo is lowered.

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2. General attenuation of the diagnostic mineral absorption bands at VNIR wavelengths. 3. Changes in the optical properties of the material. A relative increase in the spectral reflectance at increasing wavelengths, commonly termed as “spectral reddening,” has been found on Moon. (More recently, analogous “spectral bluing” has been observed on asteroid Bennu.) 4. A shift of the silicate Christiansen feature (CF) to longer wavelengths and a reduction in the strengths of the silicate Reststrahlen bands at thermal infrared (TIR) wavelengths (Lucey et al. 2017). 5. Produces the magnetic electron spin resonance of the lunar regolith. All these effects increase with soil maturity and can be quantified with various optical maturity (OMAT) parameters (not discussed here). *Interested readers could benefit from reading about another maturity parameter “Is/FeO” (Morris 1978). The change in optical properties of the lunar surface material is caused by the presence of very tiny, spherical, opaque particles of iron termed as “nanophase iron (npFe) (Cassidy and Hapke 1975; Keller and McKay 1993, 1997).” These Fe particles are typically metallic but can occur in a range of oxidation states (Thompson et al. 2016). The size and quantity of the opaque iron particles are critical in determining the optical properties of the lunar regolith affected by space weathering. Generally, the opaque iron grains comprise two distinct populations. The first is distributed throughout agglutinates and in the “inclusionrich” depositional rims on individual grains (Burgess and Stroud 2018). The second type can be found in the solar-wind-damaged amorphous rims on the host grains. The iron particles in the depositional rims are very small (~1–10 nanometers in diameter) with an average diameter of only ~3 nanometers. These small particles are responsible for reddening the spectral slope. Experimental studies have been able to demonstrate that small, nanometer-scale iron particles cause systematic reddening and darkening across the visible/near-infrared

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spectrum on the Moon. The finely dispersed larger particles (micron-scale) only lower the overall reflectance and darkens the spectra without introducing any reddening.

References Burgess KD, Stroud RM (2018) Coordinated nanoscale compositional and oxidation state measurements of lunar space-weathered material. JGR-P 123: 2022–2037 Cassidy W, Hapke B (1975) Effects of darkening processes on surfaces of airless bodies. Icar 25:371–383 Gold T (1955) The lunar surface. Mont Noti Roy Astronom Soc 115:585–604 Hawke BR, Blewett DT, Lucey PG et al (2004) The origin of lunar crater rays. Icar 170:1–16 Keller LP, McKay DS (1993) Discovery of vapor deposits in the lunar regolith. Sci 261:1305–1307 Keller LP, McKay DS (1997) The nature and origin of rims on lunar soil grains. GCA 61:2331–2341

1271 Lucey P, Korotev RL, Gillis JJ et al (2006) Understanding the lunar surface and space-Moon interactions. RiMG 60:83–219 Lucey PG, Greenhagen BT, Song E, Arnold JA et al (2017) Space weathering effects in diviner lunar radiometer multispectral infrared measurements of the lunar Christiansen feature: characteristics and mitigation. Icar 283:343–351 Morris RV (1978) The surface exposure/maturity/of lunar soils-some concepts and is/FeO compilation. In LPSC 9:2287–2297 Noble SK (2004) Turning rock into regolith: the physical and optical consequences of space weathering in the inner solar system. Brown University, Providence, RI, USA Pieters CM, Noble SK (2016) Space weathering on airless bodies. JGR-P 121:1865–1884 Thompson MS, Zega TJ, Becerra P et al (2016) The oxidation state of nanophase Fe particles in lunar soil: implications for space weathering. MAPS 51: 1082–1095

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Young Volcanism on the Moon N. Srivastava Planetary Sciences & Exploration Programme, Physical Research Laboratory, Ahmedabad, India

Introduction Our understanding of the volcanic history of the Moon in space and time is undergoing major revisions in view of studies indicating possibilities of very recent volcanism (< 100 Ma ago) in widely distributed areas across the Moon. These include a recently built volcanic field inside the Lowell crater on the farside (Srivastava et al. 2013; Gupta et al. 2014) and at least three small anomalously fresh mare patches on the nearside of the Moon (Braden et al. 2014) (Fig. 1a). Studies using global high-resolution remote sensing datasets from missions Lunar Reconnaissance Orbiter (LRO) [NASA 2009], Kaguya [JAXA 2007], and Chandrayaan-1 [ISRO 2008] have mainly contributed toward this key modification in our perception of the geological state of our nearest celestial neighbor. These missions have provided very-high-resolution panchromatic images, elevation data, hyperspectral reflectance data, and thermal emission data that are well suited for investigating the surfaces for topography, morphology, and compositional diversity at © Springer Nature Switzerland AG 2023 B. Cudnik (ed.), Encyclopedia of Lunar Science, https://doi.org/10.1007/978-3-319-14541-9

unprecedented details and to determine their model age based on crater chronology technique. Earlier, radiometric dating of lunar samples and meteorites and crater chronology-derived model ages of spectrally defined basaltic units on the Moon have indicated that the period of lunar volcanism can be broadly classified into three distinct phases: (a) early phase (> 3.8 Ga), (b) main phase (3.8–2.8 Ga), and (c) late phase (< 2.8–1.2 Ga) (e.g., Hiesinger et al. 2003; Borg et al. 2004; Terada et al. 2007; Haruyama et al. 2009; Cho et al. 2012). Inclusion of the very recent ones (