Aeronomy of Mars (Astrophysics and Space Science Library, 469) [1st ed. 2023] 9819931371, 9789819931378

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Astrophysics and Space Science Library 469

S. A. Haider

Aeronomy of Mars

Astrophysics and Space Science Library Volume 469

Series Editor Steven N. Shore, Dipartimento di Fisica “Enrico Fermi”, Università di Pisa, Pisa, Italy

The Astrophysics and Space Science Library is a series of high-level monographs and edited volumes covering a broad range of subjects in Astrophysics, Astronomy, Cosmology, and Space Science. The authors are distinguished specialists with international reputations in their fields of expertise. Each title is carefully supervised and aims to provide an in-depth understanding by offering detailed background and the results of state-of-the-art research. The subjects are placed in the broader context of related disciplines such as Engineering, Computer Science, Environmental Science, and Nuclear and Particle Physics. The ASSL series offers a reliable resource for scientific professional researchers and advanced graduate students. Series Editor: STEVEN N. SHORE, Dipartimento di Fisica “Enrico Fermi”, Università di Pisa, Pisa, Italy Advisory Board: F. BERTOLA, University of Padua, Italy C. J. CESARSKY, Commission for Atomic Energy, Saclay, France P. EHRENFREUND, Leiden University, The Netherlands O. ENGVOLD, University of Oslo, Norway E. P. J. VAN DEN HEUVEL, University of Amsterdam, The Netherlands V. M. KASPI, McGill University, Montreal, Canada J. M. E. KUIJPERS, University of Nijmegen, The Netherlands H. VAN DER LAAN, University of Utrecht, The Netherlands P. G. MURDIN, Institute of Astronomy, Cambridge, UK B. V. SOMOV, Astronomical Institute, Moscow State University, Russia R. A. SUNYAEV, Max Planck Institute for Astrophysics, Garching, Germany

S. A. Haider

Aeronomy of Mars

S. A. Haider Planetary Sciences Division Physical Research Laboratory Ahmedabad, India

ISSN 0067-0057 ISSN 2214-7985 (electronic) Astrophysics and Space Science Library ISBN 978-981-99-3137-8 ISBN 978-981-99-3138-5 (eBook) https://doi.org/10.1007/978-981-99-3138-5 © Springer Nature Singapore Pte Ltd. 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. Cover illustration: Mars on a black background isolated. Credits Shutterstock (Photo ID: 1349345675). https://www.shutterstock.com/image-photo/mars-on-black-background-isolated-red-1349345675 This Springer imprint is published by the registered company Springer Nature Singapore Pte Ltd. The registered company address is: 152 Beach Road, #21-01/04 Gateway East, Singapore 189721, Singapore

Foreword

Professor Haider has made pioneering contributions in the modeling of Planetary and Cometary Atmospheres and provided new insights and predictions that are now being confirmed. He proposed that daytime ionosphere of Mars comprises D, E and F layers as a consequence of the impact of galactic cosmic rays, X-rays and solar EUV radiations. Using an extensive range of gas phase chemistry, he investigated and modelled the cometary compounds of masses ≤ 40 amu, which explained the nine peaks of Giotto Ion Mass Spectrometer successfully. He was the first to elucidate the role of “aurora” on comets and named the VEGA spacecraft-observed energetic electrons on comet Halley as “auroral electrons”. He proposed that organic species produced by radiation-induced processes in cometary precursor grains within the solar nebula and/or in the interstellar medium could be the source of compounds with masses > 40 amu (Haider and Bhardwaj, Icarus, 177, 196–2016, 2005). This is a rare example of theoretical frame work preceding experimental verification. His recent and most innovative work on the responses of solar X-ray flares and SEP events has revolutionized the understanding of Martian auroral phenomena. This contribution has substantive implications for the coupling of SEP with ionosphere of Mars (Haider and Masoom, J. Geophys. Res, 124, 9566–9576, 2019). He has also discovered Martian magnetic storm due to arrival of CME that reached Mars after ~ 30 h of its eruption from the sun. This disturbed the E region of Martian ionosphere and enhanced the electron density significantly. He provided a new mechanism for the generation of magnetic storm on Mars (Haider et al., Space Sci. Rev., 218:33, 2022; Haider et al., Space Sci. Rev., 182:1, 2014; Haider et al., J. Geophys. Res., 117 (A5), 2012; Haider et al., Geophys. Res. Lett., 36, L1310, 2009). He has made pioneering contributions in the field of planetary atmospheres and ionospheres and their coupling with magnetospheres especially in the areas of Martian ionosphere, airglow, chemistry and aurora in the cometary coma. He specializes in development of complex models for atmospheric processes that include complex chemistry, magnetic fields and plasmas and their interaction. This has made him a unique contributor to the understanding of Martian and cometary atmospheres. Professor Haider has spent a lifetime in developing models that use basic physics, MHD and Monte Carlo method. He then demonstrated their utility towards the studies v

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Foreword

on the interaction of low and medium energy electrons with planetary atmospheres. He developed a kinetic model to estimate the current density at different potential between magnetosphere and ionosphere of Mars. The continuity-momentum model developed by him explains the various layers in the dayside and nightside ionospheres of Mars. Recently, he has pioneered an energy loss model to study galactic cosmic rays and their attenuation in the lower ionosphere of Mars (cf. Haider and Masoom, J. Geophys. Res., 124, 9566–9576, 2019); Haider et al., J. Geophys. Res., 120, 8968–8977, 2015; Haider et al., Rev. Geophys., 49, RG40001, 2011). C. T. Russell Distinguished Professor Emeritus Department of Earth, Planetary and Space Sciences University of Calfornia Los Angeles, CA, USA

Preface

Aeronomy is the science that studies upper and lower atmosphere resulting from the dissociation and ionization phenomena under the influence of the solar and particle radiation. This field has grown considerably on Mars with the launch of various missions like the Mariners, the Mars, the Vikings, Mars Global Surveyor (MGS), Mars Express (MEX), Mars Reconnaissance Orbiter (MRO) Mars Atmosphere and Volatile Evolution (MAVEN), Mangalyaan, Emirates Mars Mission (EMM) and Trace Gas Orbiter (TGO). The progress in Martian aeronomy has also been linked with the advances in ground-based measurements and modeling. The upper atmosphere is much more variable with the solar activity, solar flare and magnetic storm. The lower atmosphere of Mars is influenced from wave dynamics and dust storms. In the atmospheric study we need to know how the sun illuminates the atmosphere as a function of astronomical quantities such as the period of revolution, the rotation period, axial tilt, the inclination of the planet and the planetocentric latitude and longitude. In this book we have attempted to make a comprehensive exposition of the basic processes involved in the aeronomy of Mars. The results obtained from all aeronomical missions are discussed. We have also discussed various atmospheric and ionospheric models for Mars generated during the last four decades. The chemical reactions, rate coefficients and collision cross sections are given in Appendices for aeronomical modeling of Mars. This book is mainly divided into two parts. The first part is principally concentrated with the upper atmosphere. In this part topics on neutral atmosphere and dynamics, atmospheric escape, heating, ionospheric disturbances during low, medium and high solar activity, characteristics of Martian ionopauses, aurora and airglow are introduced. The second part is concentrated with the phenomena which occur in the middle and lower atmosphere. The topics in the middle atmosphere are concentrated on meteoroid ablation. In the lower atmosphere trace gases, gravity waves (GW), lightning, seasonal variability, thermal emissions, conductivity and dust storms are described. This book will help readers to quickly familiarize themselves in the field of Martian aeronomy. In addition it is also useful for planetary probe designers, engineers and other user’s community, e.g. planetary geologists and geophysicists.

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In this book the important models are also described. These models have been used to study the ionosphere of Mars. These models contain Boltzmann equation, continuity, momentum and energy equations, MHD and hybrid models, Monte Carlo, energy loss and meteoroid ablation methods. The photoabsorption, photoionization, electron impact and elastic cross sections are given in Appendices 1–4, respectively. The chemical reactions for the modeling of the upper ionosphere, lower ionosphere, dusty ionosphere and meteoric ions are given in Appendices 5–8, respectively. Ahmedabad, India

S. A. Haider

Acknowledgements

It has taken three long years to bring this book to its final shape. In this long journey, many of my friends have given their full support and encouragement in preparing this book. In this book I have written 26 chapters on the aeronomy of Mars. Many of the chapters appearing in this book are based on our study carried out in my 30 years long research career. I thank to publisher Springer, who always helped me and extended the dead line for the delivery of my book, whenever I requested him. Professor K. K. Mahajan, CSIR-National Physical Laboratory (NPL), spared his considerable time in reviewing the manuscript and has given very valuable inputs. All of his suggestions have helped immensely to eliminate several errors, thus enabling very smooth reading. I greatly appreciate his efforts and express my sincere thanks to him. I would like to thank Prof. Anil Bhardwaj, Director, Physical Research Laboratory (PRL), for his encouragement to write this book on Martian aeronomy. Without his support this book would not have been completed. I also acknowledge Prof. R. P. Singhal for always inspiring me to write a book on this subject. My special thanks to ISRO Chairman, who recognized my research work on Martian aeronomy and awarded me ISRO-Merit Award and released a Post-Stamp for MOM mission with my name. I am very thankful to Profs. A. K. Singhvi, J. N. Goswami, Satya Prakash, Shyam Lal and Narendra Bhandari for nominating me for the award of all three Indian Academy of Sciences (FNA, FASc and FNASc). I also acknowledge Prof. Utpal Sarkar formerly, Director Grade Scientist, who strongly recommended me for the award of the J. C. Bose Fellowship. I thank Secretary SERB for awarding me J. C. Bose National Fellowship (SB/S2/JCB-037/2016) at PRL. I especially want to thank my students/PDFs/PAs (Drs. Sharad Seth, Bhavin Pandya, Siddhi Shah, Masoom Jethwa, P. Thirupathaiah, Tikemani Bag, Vikas Singh, Disha Sawant and Mrs. Sunil Luhar, R. S. Jambusarwala and Shubham Rami) for their support during my research. Their support and enthusiasm for research have motivated me a lot. I further record my admiration for all scientific faculty, engineers and technical staff of Planetary Science Division (PSDN) for their valuable support.

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Last but not least, I thank to my beloved wife ‘Abbasi Haider’ for her understanding, unconditional support and encouragement to pursue my interest and endless patience, when it was most required. I also thank to my always positive and joyful daughter ‘Kashish Haider’, who is a powerful source of my inspiration and energy. S. A. Haider

References

Haider, S.A., Bhardwaj, A.: Radial distribution of production rates, loss rates and densities corresponding to ion masses< 40 amu in the inner coma of Comet Halley: Composition and chemistry. Icarus 177(1), 196–216 (2005) Haider, S.A., Masoom, J.: Modeling of diffuse aurora due to precipitation of H+-H and SEP electrons in the nighttime atmosphere of Mars: Monte Carlo simulation and MAVEN observation. J. Geophys. Res. Space Phys. 124(11), 9566–9576 (2019) Haider, S.A., et al.: Observations and modeling of Martian auroras. Space Sci. Rev. 218(4), 32 (2022) Haider, S.A., Mahajan, K.K.: Lower and upper ionosphere of Mars. Space Sci. Rev. 182(1), 19–84 (2014) Haider, S.A., McKenna-Lawlor, S.M.P., Fry, C.D., Jain, R., Joshipura, K.N.: Effects of solar X-ray flares in the E region ionosphere of Mars: first model results. J. Geophys. Res. Space Phys. 117(A5) (2012) Haider, S.A., Abdu, M.A., Batista, I.S., et al.: On the responses to solar X-ray flare and coronal mass ejection in the ionospheres of Mars and Earth. Geophys. Res. Lett. 36(13) (2009) Haider, S.A., Batista, I.S., Abdu, M.A., et al.: Dust storm and electron density in the equatorial D region ionosphere of Mars: Comparison with Earth’s ionosphere from rocket measurements in Brazil. J. Geophys. Res. Space Phys. 120(10), 8968–8977 (2015) Haider, S.A., Mahajan, K.K., Kallio, E.: Mars ionosphere: a review of experimental results and modeling studies. Rev. Geophys. 49(4) (2011) Haider S.A., Batista, I.S., Abdu, M.A., et al.: Flare X-ray photochemistry of the E region ionosphere of Mars. J. Geophys. Res. Space Phys. 121(7), 6870–6888 (2016) Haider S.A, Pandya, B.M., Molina-Cuberos, G.J.: Nighttime ionosphere caused by meteoroid ablation and solar wind electron-proton-hydrogen impact on Mars: MEX observation and modeling. J. Geophys. Res. Space Phys. 118(10), 6786–6794 (2013) xi

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Sheel, V., Haider, S.A.: Long-term variability of dust optical depths on Mars during MY24–MY32 and their impact on subtropical lower ionosphere: climatology, modeling, and observations. J. Geophys. Res. Space Phys. 121(8), 8038–8054 (2016) Thirupathaiah, P., Shah, S.Y., Haider, S.A.: Characteristics of solar X-ray flares and their effects on the ionosphere and human exploration to Mars: MGS radio science observations. Icarus 330, 60–74 (2019)

Contents

1

Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

1 3

2

Foundation of Ionospheric Theory . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.1 Basics of Continuity Equation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.2 Definition of Momentum Equation . . . . . . . . . . . . . . . . . . . . . . . . . . 2.3 Boundary Conditions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.4 Ambipolar Diffusion Equation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.5 Eddy Diffusion Equation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 2.6 Basics of Energy Balance . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

5 5 6 6 7 7 8 8

3

Instruments for Ionospheric Measurements on Mars . . . . . . . . . . . . . 3.1 Radio Occultation (RO) Experiment . . . . . . . . . . . . . . . . . . . . . . . . 3.2 Neutral and Ion Mass Spectrometer (NIMS) . . . . . . . . . . . . . . . . . . 3.2.1 Retarding Potential Analyzer (RPA) . . . . . . . . . . . . . . . . . 3.2.2 Langmuir Probe and Electric Field Instrument (LPEX) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 3.2.3 Mars Advanced Radar for Subsurface and Ionospheric Sounding (MARSIS) . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

11 11 12 13

Aeronomy Missions: Exploration to Mars Atmosphere . . . . . . . . . . . 4.1 Mariner 6, 7 and 9 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.2 Mars 2, 3, 4 and 5 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.3 Viking 1 and 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.4 Mars Global Surveyor (MGS) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.5 Mars Express (MEX) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.6 Mars Atmosphere and Volatile Evolution (MAVEN) . . . . . . . . . . . 4.7 Mangalyaan . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

19 19 20 21 23 24 24 27

4

15 16 17

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4.8 Emirates Mars Mission (EMM) . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4.9 ExoMars Mission . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

28 29 31

5

Thermal Structure of Mars Atmosphere . . . . . . . . . . . . . . . . . . . . . . . . . 5.1 Troposphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.2 Stratosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.3 Mesosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.4 Thermosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5.5 Exosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

37 37 38 39 39 39 40

6

Magnetic Field of Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.1 Earlier Measurements: Mariner 4, Mars 2 to 7 and Phobos 2 . . . . 6.2 Latest Measurements: MGS and MAVEN . . . . . . . . . . . . . . . . . . . . 6.3 Induced and Crustal Magnetic Fields . . . . . . . . . . . . . . . . . . . . . . . . 6.4 Mini-Magnetosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6.5 Global Modeling of Mars Magnetic Field . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

41 41 42 43 44 45 48

7

Upper Atmosphere of Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7.1 MAVEN Measurement: Neutral Compositions . . . . . . . . . . . . . . . . 7.2 Atmospheric Waves and Dynamics . . . . . . . . . . . . . . . . . . . . . . . . . 7.3 Photochemical Equilibrium on Mars . . . . . . . . . . . . . . . . . . . . . . . . 7.4 Upper Atmospheric Modeling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

49 49 50 51 52 53

8

Atmospheric Escape from Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.1 Earlier Measurements: PHOBOS-2 . . . . . . . . . . . . . . . . . . . . . . . . . 8.2 Latest Measurements: MEX and MAVEN . . . . . . . . . . . . . . . . . . . . 8.3 Thermal Escape Mechanism . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.4 Non-Thermal Escape Mechanism . . . . . . . . . . . . . . . . . . . . . . . . . . . 8.5 Modeling of Escape Flux and Density . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

57 57 58 58 59 59 60

9

Upper Ionosphere of Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.1 Photochemical Equilibrium Region . . . . . . . . . . . . . . . . . . . . . . . . . 9.2 Diffusive Equilibrium Region . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.3 Plasma Transport Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.4 Early Measurements: Mariner 4, 6, 7 and 9, Mars 4, 5 and Viking 1, 2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9.5 Latest Measurements: MGS, MEX, MARSIS and MAVEN . . . . . 9.6 Solar Zenith Angle Dependence of Mars’ Ionosphere . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

63 63 64 64 65 65 66 67

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10 Models of the Martian Ionosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.1 Boltzmann Transport Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.1.1 Continuity and Momentum Equations . . . . . . . . . . . . . . . . 10.1.2 Energy and Heat Flow Equations . . . . . . . . . . . . . . . . . . . . 10.2 Magneto-Hydrodynamic (MHD) Model . . . . . . . . . . . . . . . . . . . . . 10.3 Hybrid Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.4 Two Stream Method . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.5 Monte Carlo Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.6 AYS Approach . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.6.1 Two-Dimensional Yield Spectra . . . . . . . . . . . . . . . . . . . . 10.6.2 Three-Dimensional Yield Spectra . . . . . . . . . . . . . . . . . . . 10.6.3 Four-Dimensional Yield Spectra . . . . . . . . . . . . . . . . . . . . 10.6.4 Five-Dimensional Yield Spectra . . . . . . . . . . . . . . . . . . . . 10.7 Energy Loss Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.8 Meteoroid Ablation Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.8.1 Energy and Momentum Conservation . . . . . . . . . . . . . . . . 10.8.2 Sputtering Mass Loss . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10.8.3 Thermal Heating Mass Loss . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

69 69 70 71 72 73 74 75 75 76 76 77 78 78 79 81 81 82 84

11 Solar Flux for Ionospheric Modeling of Mars . . . . . . . . . . . . . . . . . . . . 11.1 EUV Flux Measurements: SOHO . . . . . . . . . . . . . . . . . . . . . . . . . . . 11.2 X-ray Flux Measurements: GOES . . . . . . . . . . . . . . . . . . . . . . . . . . 11.3 EUV and X-ray Flux Measurements: SORCE . . . . . . . . . . . . . . . . 11.4 EUV and X-ray Model Flux: SOLAR 2000 . . . . . . . . . . . . . . . . . . 11.5 FISM Model Flux for X-ray Flares . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

89 89 90 91 92 93 94

12 Ionization Sources of Upper Ionosphere of Mars . . . . . . . . . . . . . . . . . 97 12.1 Ionization by Solar EUV: F Region Ionosphere . . . . . . . . . . . . . . . 97 12.2 Ionization by X-rays: E Region Ionosphere . . . . . . . . . . . . . . . . . . . 98 12.3 Solar Wind Impact Ionization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 98 12.4 Dynamics of the Upper Ionosphere . . . . . . . . . . . . . . . . . . . . . . . . . 99 12.5 Chemistry of the Upper Ionosphere . . . . . . . . . . . . . . . . . . . . . . . . . 100 12.6 Modeling of the Upper Ionosphere . . . . . . . . . . . . . . . . . . . . . . . . . . 102 References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 103 13 Mars Upper Ionospheric Disturbances . . . . . . . . . . . . . . . . . . . . . . . . . . 13.1 Effects of Solar Flares on the Upper Ionosphere . . . . . . . . . . . . . . 13.2 Effects of CMEs on the Upper Ionosphere . . . . . . . . . . . . . . . . . . . 13.3 Effects of SEPs on the Upper Ionosphere . . . . . . . . . . . . . . . . . . . . 13.4 Effects of ENAs on the Upper Ionosphere . . . . . . . . . . . . . . . . . . . . 13.5 Ionospheric Modeling Due to Impact of X-ray Flares . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

107 107 108 110 111 113 113

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14 Upper Ionosphere of Mars During Low, Medium and High Solar Activity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14.1 Sunspot and Solar Activity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14.2 Effects of Solar Activity in the Upper Ionosphere . . . . . . . . . . . . . 14.3 Ionospheric Measurments: Low, Medium and High Solar Activity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14.4 Ionospheric Modeling: Low, Medium and High Solar Activity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14.5 Solar Rotation Effects on the Martian Ionosphere . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

117 117 118 119 119 120 121

15 Characteristics of Martian Ionopauses . . . . . . . . . . . . . . . . . . . . . . . . . . 15.1 Earlier Measurements of Ionopauses . . . . . . . . . . . . . . . . . . . . . . . . 15.2 Ionopause Measurements from MAVEN . . . . . . . . . . . . . . . . . . . . . 15.3 Magnetic Pile-Up Boundary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15.4 Low, Mid and High Altitude Ionopauses . . . . . . . . . . . . . . . . . . . . . 15.5 Basic Equations for Ionopause Modeling . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

125 125 126 127 127 131 133

16 Aurora and Airglow on Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.1 Discrete Aurora . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.2 Proton Aurora . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.3 Diffuse Aurora . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.4 Dayglow on Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16.5 Nightglow on Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

135 136 140 142 144 145 147

17 Middle Ionosphere of Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17.1 MGS Measurements and Meteoric Layer . . . . . . . . . . . . . . . . . . . . 17.2 MEX Measurements and Meteoric Layer . . . . . . . . . . . . . . . . . . . . 17.3 Meteoric and Atmospheric Ions . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17.4 Meteoroid Flux . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17.5 Meteors at Mars Due to Encounter of Comet C/2013 A1 . . . . . . . 17.6 Comet C/2013 A1 Observations for Meteoric Ions . . . . . . . . . . . . 17.7 Chemistry of Meteoric Ions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 17.8 Meteoric Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

151 152 153 153 155 155 156 158 158 159

18 Lower Atmosphere of Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18.1 Temperature and Pressure Measurements . . . . . . . . . . . . . . . . . . . . 18.2 Neutral Density and Composition . . . . . . . . . . . . . . . . . . . . . . . . . . . 18.3 Dynamics of the Lower Atmosphere . . . . . . . . . . . . . . . . . . . . . . . . 18.4 Atmospheric Wind . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18.5 Modeling of the Lower Atmosphere . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

161 161 162 163 163 164 167

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19 Trace Gases of Mars Atmosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.1 Ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.2 Water Vapour . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.3 Sulfur Dioxide . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.4 Methane . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19.5 Nitric Oxide . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

171 171 172 173 174 174 176

20 Seasonal Variability of Atmospheric Gases . . . . . . . . . . . . . . . . . . . . . . 20.1 Basics of Global Circulation Model (GCM) . . . . . . . . . . . . . . . . . . 20.2 Ozone Variability . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20.3 Global Mapping of Water Vapour . . . . . . . . . . . . . . . . . . . . . . . . . . . 20.4 Seasonal Variability of CO2 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20.5 Seasonal Variability of N2 , Ar, O2 and CO . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

179 179 180 181 182 182 185

21 Infrared Thermal Emissions from Mars Atmosphere . . . . . . . . . . . . . 21.1 PFS Observations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21.2 PFS Data Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21.3 Brightness Temperature of Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21.4 Planck Function and Radiance . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21.5 Radiative Transfer Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21.6 Modeling of Thermal Emission Spectra . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

187 187 188 189 192 194 195 197

22 Lower Ionosphere of Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.1 Mars 4 and 5 Measurements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22.2 Galactic Cosmic Ray (GCR) Ionization . . . . . . . . . . . . . . . . . . . . . . 22.3 Chemistry of the Lower Ionosphere . . . . . . . . . . . . . . . . . . . . . . . . . 22.4 D Layer Ionosphere of Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

199 199 200 201 203 204

23 Conductivity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23.1 Ion Conductivity in the Lower Ionosphere . . . . . . . . . . . . . . . . . . . 23.2 Effect of Dust Storm on Conductivity . . . . . . . . . . . . . . . . . . . . . . . 23.3 Ion-Dust Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23.4 Chemistry of Dusty Ionosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23.5 Atmospheric Electricity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23.6 Aerosol Charging . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

205 205 206 207 208 208 208 209

24 Dust Storms in the Lower Atmosphere of Mars . . . . . . . . . . . . . . . . . . 24.1 Infrared Dust Optical Depth: MGS Measurements . . . . . . . . . . . . 24.2 Infrared Dust Optical Depth: Mars Odyssey Observations . . . . . . 24.3 Dust Layers in the Martian Atmosphere . . . . . . . . . . . . . . . . . . . . . 24.4 Characteristics of Dust and Its Size Distributions . . . . . . . . . . . . . . 24.5 Effects of Dust in the D Region Ionosphere . . . . . . . . . . . . . . . . . .

211 211 212 214 214 214

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Contents

24.6 Density Distribution Model of Aerosol Particles . . . . . . . . . . . . . . 215 References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 217 25 Lightning on Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.1 SR Model for Lightning on Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . 25.2 Estimated SR Frequency During Nighttime . . . . . . . . . . . . . . . . . . 25.3 Search for Martian SR and Its Application . . . . . . . . . . . . . . . . . . . 25.4 SR Dependence with Height . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

221 221 225 227 228 229

26 Conclusions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 231 References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 233 Appendix: Cross Sections and Chemical Reactions for Ionospheric Modeling of Mars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 235 References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 247

About the Author

Prof. S. A. Haider is a distinguished planetary scientist of international repute. He has made pioneer contribution in the field of planetary and cometary atmospheres. Recently he has published three review papers on the Martian ionosphere in Journals Reviews of Geophysics and Space Science Reviews (Haider et al., Rev. Geophys., 49, RG4001, 2011, Haider and Mahajan, Space Sci. Rev., 182, 19–84, 2014, Haider et al., Space Sci. Rev., 218:32, 2022). These papers represent his broad understanding of planetary atmosphere-ionosphere-magnetosphere system on Mars. He has revolutionized the understanding of the effects of solar X-ray flares and CMEs on Martian ionosphere (Haider et al., J. Geophys. Res., 121, 6870–6888, 2016; Thirupathaiah et al., Icarus, 330, 60–74, 2019). Professor Haider’s achievement solves the mysteries of sudden enhancement in electron density measurements due to arrivals of solar X-ray flares and CMEs on Mars. He has developed a model and proposed for the first time that Martian ionosphere comprises D, E and F layers in the daytime ionosphere (Haider et al., J. Geophys. Res., 114, A03311, 2009, doi:1.1029/ 2008JA013709). He has also carried out an extensive study in the nighttime ionosphere of Mars. He was first to demonstrate that in Mars, the characteristic energy of magnetotail electrons was significantly different from plasma sheet region. He further explained variability of nightside Martian ionosphere due to partial screening of solar wind electrons from day side to night side along the interplanetary magnetic field lines near the terminator (Haider et al. J. Geophys. Res., 118, 1–9, 2013). The first detailed study of the Martian lower ionosphere was carried out by Prof. S. A. Haider. He proposed that Galactic cosmic ray is the main source of the lower ionosphere of Mars. He solved a detailed chemistry of the positive and negative ions in lower ionosphere and reported that dust storm reduces the densities of water cluster ions by 1–2 orders of magnitude near the surface. This was a remarkable achievement of Prof. Haider which suggested that during dust storms when optical depth changes considerably, a large hole in the ion concentrations may appear until this anomalous condition returns to the normal condition after a period of about few days (Haider et al., Rev. Geophys., 49, RG4001, 2011; Sheel and Haider, J. Geophys. Res., 121, 8038–8054, 2016). xix

xx

About the Author

He has also developed a coupled-chemistry model to calculate the densities and chemistry of masses of ions ≤ 40 amu in the coma of comet Halley. In this model he coupled ion-neutral, electron-neutral, photon-neutral and electron-ion reactions through over 600 chemical reactions. His model reproduced the nine major peaks measured by Giotto Ion Mass Spectrometer spectra throughout the inner coma. This is a rare example of theoretical frame work preceding experimental verification (Haider and Bhardwaj, Icarus, 177, 196–216, 2005). Professor Haider has made pioneer contributions in the modelling of planetary atmospheres, ionospheres and their coupling with magnetospheres especially in the areas of Martian ionosphere, airglow, chemistry and aurora in cometary coma. He has discovered magnetic storms and their possible mechanisms due to arrival of CME on Mars. Based on his innovative contribution in planetary sciences he was elected as a fellow of all three Indian Academy of Sciences and the J. C. Bose National fellow. He has also authored and edited two books on the Modelling of Planetary Atmospheres and Advances in Geosciences. He has been elected as the president for Planetary Sciences 2012–2014 by Asia Oceania Geosciences Society. Recently, he has been honoured from ISRO-Merit Award-2017.

Acronyms

ACS ADU AIS ASPERA AYS CaSSIS CB CEM CME CRISM CSPA DSMC DSN ELF EM EMIRS EMM EMUS EUV EXI FISM FREND GCM GCMS GCR GCV GDS GOES GW HRSC IMF

Atmospheric Chemistry Suite Analogue Digital Units Active Ionospheric Sounding Analyser of Space Plasmas and Energetic Atoms Analytical Yield Spectrum Colour and Stereo Surface Imaging System Cavity Boundary Channel Electron Multiplier Coronal Mass Ejections Compact Reconnaissance Imaging Spectrometer Charge-Sensitive Pre-amplifier Direct Simulation Monte Carlo Deep Station Network Extremely Low Frequency Electro-magnetic Emirates Infrared Spectrometer Emirates Mars Mission Emirates Mars Ultraviolet Spectrograph Extreme Ultraviolet Emirates Exploration Imager Flare Irradiance Spectral Model Fine Resolution Epithermal Neutron Detector Global Circulation Models Gas Chromatograph Mass Spectrometer Galactic Cosmic Rays Generalized Coefficient of Variation Global Dust Storm Geostationary Operational Environmental Satellite Gravity Waves High-Resolution Stereo Camera Interplanetary Magnetic Field xxi

xxii

IUVS LAP LEMA LMD LPEFI LPWA Ls LW MAG/ER MaRs MARSIS MAVEN MAWD MCC MCD MENCA MEX MGCM MGITM MGS MHD MNP MOC MOLA MPB MRO MSE MSL MSM MTGCM MY NGIMS NIMS NOMAD OMEGA PFS QMA RO ROSE RPA S2K SEM SEP SH SIM

Acronyms

Imaging Ultraviolet Spectrometer Lyman Alpha Photometer Lightning Experiment for Mars Laboratoire de Météorologie Dynamique Langmuir Probe and Electric Field Instrument Langmuir Probe and Waves Antenna Solar Longitude Long wavelength Magnetometer and Electron Reflectometer Mars Radio Science Mars Advanced Radar for Subsurface and Ionospheric Sounding Mars Atmosphere and Volatile Evolution Mars Atmospheric Water Detector Mars Colour Camera Mars Climate Database Mars Exospheric Neutral Composition Analyser Mars Express Mars Global Circulation Model Mars Global Ionosphere Thermosphere Model Mars Global Surveyor Magnetohydrodynamics Mars Nearest Point Mars Orbiter Camera Mars Orbiter Laser Altimeter Magnetic Pile-Up Boundary Mars Reconnaissance Orbiter Mars-Sun-Electric field Mars Science Laboratory Methane Sensor for Mars Mars Thermosphere Global Circulation Model Martian Years Neutral Gas and Ion Mass Spectrometer Neutral and Ion Mass Spectrometer Nadir and Occultation for Mars Discovery Observatoire Pour La Mineralogie I’Eau Las Glaces et’Active Planetary Fourier Spectrometer Quadrupole Mass Analyser Radio Occultation Radio Occultation Science Experiment Retarding Potential Analyser SOLAR2000 model Solar EUV Monitor Solar Energetic Particle Spherical Harmonic Spectral Irradiance Monitor

Acronyms

SIP SOHO SOLSTICE SORCE SPICAM SR SSI STATIC SW SWEA SWIA SWPC SZA TAUS TEC TES TGO THEMIS TID TIS TSI USO UVD VLF XPS XRF

xxiii

Solar Irradiance Platform Solar and Heliospheric Observatory SOLar STellar Irradiance Comparison Experiment Solar Radiation and Climate Experiment Spectroscopy for Investigation of Characteristics of the Atmosphere of Mars Schuman Resonance Solar Spectral Irradiance Supra Thermal and Thermal Ion Composition Short Wavelength Solar Wind Electron Analyser Solar Wind Ion Analyser Space Weather Prediction Centre Solar Zenith Angle Solar Wind Plasma Experiment Total Electron Content Thermal Emission Spectrometer Trace Gas Orbiter Thermal Emission Imaging System Travelling Ionospheric Disturbances Thermal Infrared Imaging Spectrometer Total Solar Irradiance Ultrastable Oscillator Ultraviolet Doublet Very Low Frequency X-ray UV Photometer System X-ray Fluoroscence

Chapter 1

Introduction

The atmosphere of Mars is about 100 times thinner than the Earth’s atmosphere. It has 95% carbon dioxide. The average temperature of Mars is about 220 °K. It varies from equator to poles and also with seasons. The dust storms of Mars are the largest in the solar system. It usually occurs when Mars is near its aphelion. Mars might had once harboured life. A number of researchers have argued that oceans may have covered the surface of Mars in the past, providing an environment for life to develop (Mckay 1997; Farmer 1998; Baker et al. 1991; Cabrol and Grin 2005). Mars is cold desert today. The liquid water may be present underground, providing a potential refuge for any life that might exist there. Several studies have shown that there is abundant water ice beneath the surface. Mars experiences all four seasons (Barth 1974; Klein 1979; Haberle et al. 2003) that the Earth does (see Fig. 1.1). Table 1.1 shows the range of solar longitudes to each Martian season in northern and southern hemispheres. Since the year is longer on Mars, its seasons are not of the same length as on Earth. Mars also contains ionosphere, which is composed of ions and electrons. It contains lower and upper ionosphere, which are formed by Galactic Cosmic Rays (GCR) and solar Extreme Ultraviolet (EUV)/X-ray radiation respectively (Haider et al. 2007, 2009a, b, 2011; Pandya and Haider, 2014; Shah et al. 2021). Mars does not have dipole magnetic field. In the upper atmosphere the magnetic field was measured by a magnetometer onboard MGS and MAVEN (Acuña et al. 1998; Connerney et al. 1999; Acuña et al. 2001; Connerney et al. 2015a, b; Mittelholz et al. 2018). The magnetic field on Mars is very weak, which does not significantly contribute to the solar wind-Mars interaction. The important physical parameters (axial tilt, eccentricity, gravity, diameter, rotation, escape velocity, distance from Sun, mean temperature and atmospheric pressure) of Mars and Earth are given in Table 1.2 for comparison. The atmosphere of Mars causes a great number of obstacles for human exploration of this planet. It prevents from the liquid water on the surface and allows higher levels of radiation than Earth, which humans can barely tolerate and it becomes very difficult to grow food even in the green house. Mars has acquired its atmosphere from the interior. Most of this outgasing took place during billion years ago. Once the planet cooled, it released little gas. There is evidence that in the past a denser Martian © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_1

1

2

1 Introduction

Fig. 1.1 Schematic representation of Martian seasons (https://mars.nasa.gov/resources/7889/sea sons-in-the-martian-year-as-the-red-planet-orbits-the-sun/) Table. 1.1 Martian Season

Table. 1.2 Atmospheric Parameters for Earth and Mars

Solar longitude (Ls)

Northern hemisphere

Southern hemisphere

0°–89°

Spring

Autumn

90°–179°

Summer

Winter

180°–269°

Autumn

Spring

270°–359°

Winter

Summer

Earth

Physical parameters

Mars

Radius (km)

6378

3396

Mass (kg)

5.97 × 1024

6.39 × 1023

1

1.52

Average distance from Sun (AU) (Wm−2 )

1367

589

Orbital eccentricity

0.017

0.093

Length of year (Mars sols)

355.6

668.6

Mean solar constant

Length of year (Earth days)

365.25

687.0

Spin-axis inclination (°)

23.44

25.19

Rotation frequency (s−1 )

7.29 × 10–5

7.09 × 10–5

Solar day (h)

24.00

24.66

Surface gravity (ms−2 )

9.80

3.71

Average surface pressure (hPa)

1013

6.1

Average surface temperature (K)

220

210

References

3

atmosphere may have warmed the planet enough to allow water to flow on the surface. Physical features closely resembling shorelines, gorges, riverbeds and islands suggest that great rivers once flowed on the planet. One of the greatest mysteries about Mars is the loss of its atmosphere and water, in particular, the loss of oxygen. There is evidence of significant loss to space in the ratios of isotopic abundances of stable elements, for example, D/H, 13 C/12 C, 15 N/14 N, and 38 Ar/36 Ar. The atmosphere of Mars is also in thermodynamic equilibrium as the atmospher of Earth. Therefore, Mars atmosphere absorbs exactly the same amount of thermal energy that it radiates into space. In this process, absorbed energy creates the atmospheric dynamics of wind and weather. All of this energy is eventually degraded into thermal radiation at infrared wavelengths and radiated into space.

References Acuña, M.H., et al.: Magnetic field and plasma observations at Mars: initial results of the Mars Global Surveyor mission. Science 279(5357), 1676–1680 (1998) Acuña, M.H., et al.: Magnetic field of Mars: summary of results from the aerobraking and mapping orbits. J. Geophys. Res. Planets 106(E10), 23403–23417 (2001) Baker, V.R., Strom, R.G., Gulick, V.C., et al.: Ancient oceans, ice sheets and the hydrological cycle on Mars. Nature 352(6336), 589–594 (1991) Barth, C.A.: The atmosphere of Mars. Annu. Rev. Earth. Planet Sci. 2, 333 (1974) Cabrol, N.A., Grin, E.A.: 10 Ancient and recent lakes on Mars. In: Water on Mars and life, pp. 235– 259. Springer, Berlin, Heidelberg (2005) Connerney, J.E.P., Acuna, M.H., Wasilewski, P.J.P.A., et al.: Magnetic lineations in the ancient crust of Mars. Science 284(5415), 794–798 (1999) Connerney, J.E.P., Espley, J., Lawton, P., et al.: The MAVEN magnetic field investigation. Space Sci. Rev. 195(1), 257–291 (2015a) Connerney, J.E., Espley, J.R., DiBraccio, G.A., et al.: First results of the MAVEN magnetic field investigation. Geophys. Res. Lett. 42(21), 8819–8827 (2015b) Farmer, J.: Thermophiles, early biosphere evolution, and the origin of life on Earth: implications for the exobiological exploration of Mars. J. Geophys. Res.: Planets 103(E12), 28457–28461 (1998) Haberle, R.M., Murphy, J.R., Schaeffer, J.: Orbital change experiments with a Mars general circulation model. Icarus 1; 161(1), 66–89 (2003) Haider, S.A., Singh, V., Choksi, V.R., Maguire, W.C., Verigin, M.I.: Calculated densities of H3 O+ (H2 O)n, NO2 − (H2 O)n, CO3 − (H2 O)n and electron in the nighttime ionosphere of Mars: Impact of solar wind electron and galactic cosmic rays. J. Geophys. Res. Space Phys. 112(A12) (2007) Haider, S.A., Abdu, M.A., Batista, I.S., et al.: D, E, and F layers in the daytime at high-latitude terminator ionosphere of Mars: comparison with Earth’s ionosphere using COSMIC data. J. Geophys. Res. Space Phys. 114(A3) (2009a) Haider, S.A., Abdu, M.A., Batista, I.S., et al.: On the responses to solar X-ray flare and coronal mass ejection in the ionospheres of Mars and Earth. Geophys. Res. Lett. 36(13) Haider, S.A., Mahajan, K.K., Kallio, E.: Mars ionosphere: a review of experimental results and modeling studies. Rev. Geophys. 49(4) (2011) Klein, H.P.: The Viking mission and the search for life on Mars. Rev. Geophys. 17(7), 1655–1662 (1979) Mckay, C.P.: The search for life on Mars. In: Planetary and interstellar processes relevant to the origins of life, pp. 263–289 (1997)

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1 Introduction

Mittelholz, A., Johnson, C.L., Morschhauser, A.: A new magnetic field activity proxy for Mars from MAVEN data. Geophys. Res. Lett. 45(12), 5899–5907 (2018) Pandya, B.M., Haider, S.A.: Numerical simulation of the effects of meteoroid ablation and solar EUV/X-ray radiation in the dayside ionosphere of Mars: MGS/MEX observations. J. Geophys. Res. Space Phys. 119(11), 9228–9245 (2014) Shah, S.Y., et al.: A coupled model of the D and E regions of Mars’ ionosphere for flare and non-flare electron density profiles. Icarus 361, 114403 (2021)

Chapter 2

Foundation of Ionospheric Theory

Abstract In this chapter we have described the basic theory of ionosphere such as continuity equation, momentum equation, ionospheric boundary conditions, ambipolar diffusion, eddy diffusion and thermal energy balance. Using these equations the ion densities, electron densities and temperature structures can be studied in the lower and upper ionosphere of Mars. Keywords Ionospheric tools · Basic equations · Photochemical region

2.1 Basics of Continuity Equation Continuity equation describes the transport of a quantity. It requires source and sink terms, which are often, but not always conserved such as the density of a molecular species which can be created or destroyed by chemical reactions. Continuity equations are represented in terms of convection, diffusion equation, Boltzmann transport equation and Nover-Stokes equation. The distribution of the charged particles (electron and ions) in the Martian ionosphere is governed by continuity equation (Shinagawa and Cravens 1989; Shinagawa and Bougher 1999; Ma et al. 2004). Continuity equation calculates the production rates, loss rates and diffusion in the Martian ionosphere. In 1-D continuity equation the plasma flows in vertical direction Z. In 2-D continuity equation the plasma flows in Y and Z directions, where Y is a horizontal direction. In 3-D continuity equation the plasma flows in X, Y, and Z directions, where X is perpendicular to Y and Z directions. For steady state condition the production and loss rates are equal. The vertical diffusion can be neglected under steady state equilibrium condition. Then ion and electron density can be calculated by iteration process. In the ionosphere of Mars, the chemical life time (τ = l−1 ) is much less up to 200 km than the molecular diffusion time (τ m = H 2 /D); where l is loss coefficient, H and D are scale height and diffusion coefficient respectively (St Maurice and Schunk 1977).

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2.2 Definition of Momentum Equation The continuity equation in the fluid dynamics describes that in the steady state process, speed at which mass leaves the system is adequate to the speed at which mass enters a system. The differential form of the continuity equation is given as ∂n + ∇ · (n · u→) = 0 ∂t

(2.1)

where t is time, n is fluid density and u→ is a diffusion velocity vector field. Under = 0, then steady state condition, ∂n ∂t ∇ · (n · u→) = 0 n.→ u = constant This is known as diffusive equilibrium condition, where flux is constant. This condition is valid where chemical reactions are not very important and the plasma reached at stable distribution due to ambipolar diffusion. In this region the collision is negligible and plasma can escape from the exosphere. The diffusion velocity u→ can be represented in three dimensions u→x , u→y , and u→z . Consider a fluid of lengths dx, dy, and dz in X, Y, and Z directions respectively. The second term of Eq. (2.1) represents derivatives of diffusion flux (ambipolar diffusion and eddy diffusion flux).

2.3 Boundary Conditions In order to solve the above equation, lower and upper boundary conditions are required: (1) photochemical equilibrium condition is imposed on the ion densities at the lower boundary, where the vertical velocity u→z is kept zero, and (2) at the upper boundary the vertical velocity is again set to be zero at th ionopause altitude. It should be noted that ionopause altitudes are changing with solar wind dynamic pressure. →

∂ u

In presence of spatial derivative of horizontal plasma flow ∂ yy in Eq. (2.1), it is necessary to specify the boundary condition for u→y . Shinagawa and Cravens (1989) used Newtonian distribution to obtain horizontal divergence term, where magnetic field pressure varies as Cos2 χ in the horizontal direction (where χ is a SZA, which is zero at the sub-solar point). This equation represents the approximate ion loss term due to the divergence of the horizontal velocity in the continuity Eq. (2.1). This loss term reduced the electron density by ~70% near Mars’ ionopause in presence of horizontal magnetic field of ~30–40 nT.

2.5 Eddy Diffusion Equation

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2.4 Ambipolar Diffusion Equation In planetary ionosphere, electrons, ions and minor constituents diffuse through ambient atmosphere. The electrons diffuse faster than the ions. A polarization field exists between ions and electrons, which prevents charge separation. In presence of polarization electric field E, the electrons and ions diffuse with the same velocity. This process is called ambipolar diffusion. The ambipolar diffusion is dominant in the upper atmosphere. The plasma diffusion velocity u→d for steady state condition can be represented as follows: [ u→d = − Da

1 ∂N 1 (1 + α) ∂ T + + N ∂z Hp T ∂z

] (2.2)

Hp =

k (T e + T i ) mi g

(2.3)

Da =

k (T e + T i ) m i νin

(2.4)

where N is the density, H p is plasma scale height, g is gravity, mi is mass of ith ion, k is Boltzmann constant, α is thermal diffusion factor, T is temperature, ν in is ion-neutral frequency, T i and T e are the ion and electron temperatures respectively. The ion-neutral collision frequency can be represented as: ( νin = 2π

α e2 μin

) 21 n

(2.5)

where α is the polarizability, e is electronic charge, μin = mi mn /(mi + mn ) is the reduced mass and n is number density of the neutrals. The ion neutral frequency is more important than the electron neutral frequency because mi >> me . The mn and me are neutral and electron masses respectively. Da is the ambipolar diffusion coefficient. For T e = T i , Da = 2 Di , which is twice the ion diffusion coefficient Di .

2.5 Eddy Diffusion Equation The eddy diffusion is a process by which substances are mixed. It is dominant in the lower atmosphere. The eddy diffusion velocity u→d can be written as given by Eq. (2.2) [ u→d = − K d

1 1 ∂T 1 ∂N + + N ∂z H T ∂z

] (2.6)

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where K d is eddy diffusion coefficient and H is neutral scale height. In eddy diffusion coefficient, thermal diffusion factor is zero. Under steady state condition the diffusion velocity can be described by the sum of Eqs. (2.2) and (2.6). The Eq. (2.2) represents the ambipolar diffusion term, while eddy diffusion term is represented by Eq. (2.6). The eddy diffusion is important below the turbopause, while molecular diffusion is predominant above the turbopause.

2.6 Basics of Energy Balance The temperature profiles of ions and electrons in the planetary ionosphere can be estimated from heat balance equation in the form of kinetic energy, thermal conductivity, heating and cooling rates. If the ionosphere is pervaded by the magnetic field, the conductivity will be strongly controlled by the direction of this field. The initial source of energy for heating the charged particles in the ionosphere is obtained from photoionization process. In this process energetic photoelectrons are created, which can lose energy due to elastic and inelastic collisions with the neutral species leading to rotational, vibrational or electronic excitation of neutrals. The photoelectrons are highly energetic ranging between 1 and 100 eV (Nagy et al. 1980). If the energy of primary photoelectron is more than the ionization potential, secondary ionization may occur. When energy is less than the ionization potential, the excess energy is then lost by excitation of atmospheric constituents. The thermal electrons cool through (1) elastic and inelastic collisions with neutral particles, (2) elastic collisions with ions, (3) rotational and vibration excitation of CO2 and CO, (4) excitation of fine structure levels of atomic oxygen by electron impact, and (5) electronic excitation of O to 1 D level from the 3 P ground level. The most important of these are processes # (2), (4), and (5). In the planetary ionosphere the electron temperature is larger than the ion and neutral temperatures. The thermal conductivity of electron is higher than the ion conductivity (Haider et al. 2010).

References Haider, S.A., Seth, S.P., Brain, D.A., et al.: Modeling photoelectron transport in the Martian ionosphere at Olympus Mons and Syrtis Major: MGS observations. J. Geophys. Res. Space Phys. 115(A8) (2010) Ma, Y., Nagy, A.F., Sokolov, I.V., Hansen, K.C.: Three-dimensional, multispecies, high spatial resolution MHD studies of the solar wind interaction with Mars. J. Geophys. Res. Space Phys. 109(A7) (2004) Maurice, J.P., Schunk, R.W.: Diffusion and heat flow equations for the mid-latitude topside ionosphere. Planet Space Sci 25(10), 907–920 (1977)

References

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Nagy, A.F., Cravens, T.E., Smith, S.G., et al.: Model calculations of the dayside ionosphere of Venus: ionic composition. J. Geophys. Res. Space Phys. 85(A13), 7795–7801 (1980) Shinagawa, H., Bougher, S.W.: A two-dimensional MHD model of the solar wind interaction with Mars. Earth Planets Space 51(1), 55–60 (1999) Shinagawa, H., Cravens, T.E.: A one-dimensional multispecies magnetohydrodynamic model of the dayside ionosphere of Mars. J. Geophys. Res. Space. Phys. 94(A6), 6506–6516 (1989)

Chapter 3

Instruments for Ionospheric Measurements on Mars

Abstract The most important instruments, which have been used extensively to the Mars’ ionosphere are: (1) Radio Occultation (RO), (2) Neutral and Ion Mass Spectrometer (NIMS), (3) Mars Exospheric Neutral Composition Analyser (MENCA) (4) Retarding Potential Analyzer (RPA), (5) Langmuir Probe and Electric Field Instrument (LPEFI), and (6) Mars Advanced Radar for Subsurface and Ionospheric Sounding (MARSIS). In brief these instruments are described below. Keywords Ionospheric instruments · Aeronomy instruments · Measurements

3.1 Radio Occultation (RO) Experiment The RO is a powerful instrument to measure the electron density profile in the ionosphere of Mars. When the spacecraft passes behind the planet, radio communication passing through the Martian atmosphere experiences a bending leading to Doppler shift, which is recorded as phase information at the Deep Station Network (DSN) located on Earth. Schematic representation of ray diagram of RO method is shown in Fig. 3.1. This Doppler residual can be used to study both the neutral atmosphere and ionosphere parameters. First the bending angle and then the refractivity are derived from the Doppler residuals. For neutral atmosphere, the refractivity is greater than zero, whereas for the ionosphere it is less than zero. In RO experiment S and X band radio waves are used simultaneously to know the frequency dependence of phase and amplitude variation. S band is sensitive to 2.3 GHz frequency to measure the plasma density and X band is sensitive to 8.4 GHz frequency to measure the neutral density of planetary atmospheres. To develop this instrument we need transponders and amplifiers, whose noise level should be very low. X band and S band downlink will be transmitted simultaneously and coherently by transponder. The frequency of onboard oscillator needs to be stable enough, which will be received by ground station. To stabilize both link frequencies, an Ultra-Stable Oscillator (USO) is required. Changes in the transmitted signal will be attributed to the changes due to Martian atmosphere. It can be useful to do one-way (mono-static) observation and operation of both downlink frequencies for sounding © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_3

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Fig. 3.1 Schematic diagram of ray path, asymptotic angle and DSN receiving antenna of RO experiment (from Imamura et al. 2017)

planetary atmospheres. One way will cover the ingress and egress of the occultation. The two-way (by-static) observation is limited to ingress of occultation only.

3.2 Neutral and Ion Mass Spectrometer (NIMS) Figure 3.2 represents schematic diagram of NIMS, which measures both neutral gas molecules and charged ions (Waite et al. 2004). In neutral particle mode, a high positive voltage is applied on the entrance grid to avoid the entrance of the positive charged ions to the spectrometer. In this mode, only neutral gas molecules will be entering inside the instrument, which will be positively ionized using electron impact ionization. The positive ions will be accelerated towards the quadrupole mass filter, where ions will be filtered as per their m/q ratio. In the charged particle mode, the positive charged ions will be allowed to enter into the instrument by applying ground/negative voltage on entrance grid. The positive ions will be accelerated towards the mass filter, where ions will be filtered as per their m/q ratio. In this mode electrons will have acceleration in the reverse direction and hence will not be detected. The ions will be detected using Channel Electron Multiplier (CEM) detector. The output signal of the CEM will be connected to high speed charge sensitive pre-amplifier (CSPA). The output of the CSPA will be digitized using high speed ADC and will be interfaced to FPGA. The FPGA will communicate with the TM/TC/BDH for the payload operation and data communication. The processing electronics will be developed using FPGA, which will monitor all the operations and acquire the data from ADC. Raw power of the satellite is converted to high voltage and low voltage supply using DC/DC converters, and will be utilized to operate the instrument.

3.2 Neutral and Ion Mass Spectrometer (NIMS)

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Fig. 3.2 Schematic representation of NIMS for ions and neutral species measurements in the planetary ionosphere (from Waite et al. 2004)

The MENCA onboard Magalyaan is a Quadrupole Mass Analyser (QMA)-based instrument with a cylindrical sensor. This instrument is shown in Fig. 3.3. In this figure cylinder consists of an ionizer section, QMA and a detector assembly. The sensor is attached with the electronics (rectangular box) that controls the instrument and telemeters the data through the spacecraft. This instrument is measuring the compositions of Martian neutral exosphere in the 1–300 amu mass range in the equatorial and low latitudes of Mars (Bhardwaj et al. 2016).

3.2.1 Retarding Potential Analyzer (RPA) The RPA is one of the powerful techniques to make in situ measurements of ion densities and their temperatures. This instrument consists of four perforated electrodes (grids) and a solid electrode (collector) that are biased to various voltages to accomplish energy-based ion filtering as shown in Fig. 3.4. The floating grid establishes a floating potential (Vf ) that traps the plasma outside the sensor and minimizes probe perturbation on the plasma. The first electron repelling grid biased at a negative voltage (Ve-) removes the electrons from the particle beam. The ion-retarding grid biased at a variable potential (Vion ) progressively shields more energetic ions as it is swept from a low potential to high potential; the ion current is intercepted by the inner most electrode, i.e. the collector, after passing through a second electron-repelling grid that prevents secondary electron emission from the collector.

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Fig. 3.3 The flight model of MENCA instrument onboard Mangalyaan (Credit ISRO)

Fig. 3.4 Schematic representation of typical structure and operations of ion RPA instrument (from Heubel et al. 2014)

3.2 Neutral and Ion Mass Spectrometer (NIMS)

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3.2.2 Langmuir Probe and Electric Field Instrument (LPEX) The LPEX is based on current and voltage measurements, which is shown in Fig. 3.5. In this figure two booms are shown for electric field measurement. The Langmuir Probe exists in the middle of the diagram. Both booms are separated by a distance (2.5 m), which enables the probe to be adequately far from the central payload to reduce the field interface caused by the payload structure and also to satisfy the Debye length condition. The length of the boom should be more than the Debye length of plasma. The electric fields can be measured from few mV/m to few V/ m. Current–voltage are collected using Langmuir probe/electric field sensors, which are pre-amplified then transmitted to the central electronics through harness running along the boom. This experiment onboard MAVEN is measuring the electron density and electric field in the ionosphere, bow shock and ion tail regions of Mars in situ along the orbit of the spacecraft. In these regions electron density and electric field are changing significantly due to highly dynamical processes.

Fig. 3.5 Schematic representations of Langmuir Probe and Electric Field Instrument (LPEX) (from Durga Prasad et al. 43rd COSPAR Assembly, 28 January–4 February, 2021) (Abstract # C3.2-003521)

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3.2.3 Mars Advanced Radar for Subsurface and Ionospheric Sounding (MARSIS) The MARSIS is a low to high frequency (0.1–5.5 MHz) ionospheric sounding radar onboard MEX, which has been sounding the planet since 2005. It is working in two different operational modes. One of these is the Active Ionospheric Sounding (AIS) (Picardi et al. 2004; Gurnett et al. 2005), which, basically, is a topside sounding mode to measure the topside electron density profile. This profile can then be used to obtain the Total Electron Content (TEC) of the topside ionosphere. The other operational mode is the subsurface mode in which the radar sounds the surface and subsurface of the planet to identify and to measure the material below it. In the subsurface mode, TEC is derived as a by-product from the analysis of signal distortion caused by the dispersive ionosphere (Safaenili et al. 2007; Leblanc et al. 2008a, b; Mouginot et al. 2008; Cartacci et al. 2013; Sánchez-Cano, et al. 2015). In AIS mode MARSIS measures reflections from normal ionosphere in nadir direction (called vertical echoes) and ionization bulges in the oblique directions (called oblique echoes) (Andrews et al. 2014; Dieval et al. 2018; Venkateswara Rao et al. 2019; Duru et al. 2006). The peak frequencies of these echoes are generally the same. However, in some cases frequencies of oblique echoes were much larger than those of vertical echoes (Venkateswara Rao et al. 2019). A typical MARSIS ionogram is shown in Fig. 3.6a, b. The vertical lines at the top left corner of this figure represent the harmonics of electron plasma oscillations. This figure also depicts ionospheric and surface echoes. The traces of these echoes can be inverted to get an electron density profile and TEC respectively. The horizontal lines appearing at the low-frequency end correspond to the electron cyclotron echoes. The separation between these lines is used to derive the intensity of the total magnetic field local to the spacecraft (Akalin et al. 2010). As mentioned above, when MARSIS is passing through an ionization bulge, a second ionospheric echo trace, called oblique echo trace, is also observed as shown in Fig. 3.6a.

References

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Fig. 3.6 Two ionograms observed on 14 August (a), and 10 August (b) 2005 are illustrating various ionospheric features generally observed by MARSIS. Electron plasma oscillation harmonics and ionospheric echos can be seen in both ionograms. A strong surface reflection and oblique echo are seen in ionogram 3.6a but it was not found in ionogram 3.6b (from Gurnett et al. 2005)

References Akalin, F., Morgan, D.D., Gurnett, D.A., et al.: Dayside induced magnetic field in the ionosphere of Mars. Icarus 206(1), 104–111 (2010) Andrews, J., André, D., Opgenoorth, M., et al.: Oblique reflections in the Mars Express MARSIS data set: stable density structures in the Martian ionosphere. J. Geophys. Res. Space Phys. 119(5), 3944–3960 (2014)

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Bhardwaj, A., Thampi, S.V., Das, T.P., et al.: On the evening time exosphere of Mars: result from MENCA aboard Mars Orbiter Mission. Geophys. Res. Lett. 16; 43(5), 1862–1867 Cartacci, M., Amata, E., Cicchetti, A., et al.: Mars ionosphere total electron content analysis from MARSIS subsurface data. Icarus 223(1), 423–437 (2013) Diéval, C., Kopf, A.J., Wild, J.A.: Shapes of magnetically controlled electron density structures in the dayside Martian ionosphere. J. Geophys. Res. Space Phys. 123(5), 3919–3942 (2018) Duru, F., Gurnett, D.A., Averkamp, T.F., et al.: Magnetically controlled structures in the ionosphere of Mars. J. Geophys. Res. Space Phys. 111(A12) (2006) Gurnett, D.A., Kirchner, D.L., Huff, R.L., et al.: Radar soundings of the ionosphere of Mars. Science 310(5756), 1929–1933 (2005) Heubel E V (2014) Enhancing retarding potential analyzer energy measurements with micro-aligned electrodes (Doctoral dissertation, Massachusetts Institute of Technology). Imamura, T., Ando, H., Tellmann, S., et al.: Initial performance of the radio occultation experiment in the Venus orbiter mission Akatsuki. Earth Planets Space 69(1), 1–11 (2017) Leblanc, F., Witasse, O., Lilensten, J., et al.: Observations of aurorae by SPICAM ultraviolet spectrograph on board Mars Express: Simultaneous ASPERA-3 and MARSIS measurements. J. Geophys. Res. Space. Phys. 113(A8) (2008a) Leblanc, F., Aplin, K., Yair, Y., et al.: Planetary atmospheric electricity, vol. 30. Springer (2008b) Mouginot, J., Kofman, W., Safaeinili, A., Hérique, A.: Correction of the ionospheric distortion on the MARSIS surface sounding echoes. Planet Space Sci. 56(7), 917–926 (2008) Picardi, G., Biccari, D., Seu, R.: MARSIS: Mars advanced radar for subsurface and ionosphere sounding In: Mars Express: The Scientific Payload, vol. 1240, pp. 51–69 (2004) Safaeinili, A., Kofman, W., Mouginot, J., et al.: Estimation of the total electron content of the Martian ionosphere using radar sounder surface echoes. Geophys. Res. Lett. 34(23) (2007) Sánchez-Cano, B., Morgan, D.D., Witasse, O., et al.: Total electron content in the Martian atmosphere: a critical assessment of the Mars Express MARSIS data sets. J. Geophys. Res. Space Phys. 120(3), 2166–2182 (2015) Venkateswara Rao, N., Leelavathi, V., Mohanamanasa et al.: Enhanced ionization in magnetic anomaly regions of the Martian lower ionosphere associated with dust storms. J. Geophys. Res. Space Phys. 124(4), 3007–3020 (2019) Waite, J.H., Lewis, W.S., Kasprzak, W.T., et al.: The Cassini ion and neutral mass spectrometer (INMS) investigation. Space Sci. Rev. 114(1), 113–231 (2004)

Chapter 4

Aeronomy Missions: Exploration to Mars Atmosphere

Abstract Mars’ upper atmosphere has been observed by Mariners 6, 7, 9; Mars 2, 3, 4, 5; Viking 1 and 2, MGS, MEX, MAVEN, Mangalyaan, EMM, and ExoMars but there have been very few measurements in the lower atmosphere of Mars. Brief descriptions of these missions are given below. Keywords Aeronomy missions · Mars exploration · Atmosphere

4.1 Mariner 6, 7 and 9 Mariner 6 and 7 were launched on February 24 and March 27, which flew by Mars on July 31 and August 5, 1969, respectively. They carried four atmospheric instruments (1) RO Experiment (Kliore et al. 1972), (2) Infrared Spectrometer (Hanel et al. 1972), (3) Radio meter (Kieffer et al. 1973), and (4) Ultraviolet Spectrometer (Barth et al. 1972). The exploration of Mars with an orbiter was started by Mariner 9, which was launched on May 30, 1971. It also carried same atmospheric instruments, which were aboard on Mariner 6 and 7 spacecrafts. During the first 40 days, a total of 160 occultation measurements were performed from Mariner 9. These data corresponded to latitudes between 34 and 65° N and solar zenith angle varied from 105° to 57°. Another set of occultation measurements was obtained during May–June, 1972. The latitude coverage ranged from 86° N to 80° S with solar zenith angle varying between 70° and 100°. It also observed a global dust storm. In Fig. 4.1, Mariner 9 orbiter is shown with four aeronomy instruments flown at Mars (Parks 1973). The spacecraft sensors, solar panel, antennas and other necessary components for orbiting Mariner 9 are also shown in this figure.

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Fig. 4.1 Mariner 9 spacraft with four aeronomy instruments flown at Mars from NASA (from Parks 1973)

4.2 Mars 2, 3, 4 and 5 Mars 2 and 3 consisted of identical spacecraft, each with an orbiter and a Lander. These spacecrafts were launched on May 19, 1971 and May 28, 1971 respectively and completed their missions by August 22, 1972. For the first time these missions have observed ion escape rates of the order of 1024 –1025 s−1 from the Mars’ atmosphere, which correspond to evacuation of the total oxygen content in the present day atmosphere of Mars in less than 108 –109 years (Vaisberg and Smirnov 1986). Mars 4 and Mars 5 were launched on July 21 and 25, 1973 respectively. First radio sounding of the Martian atmosphere above the dark surface of the planet were carried out during the exits from behind the planet by the spacecrafts Mars 4 and Mars 5 on 10 and 18 February, 1974, respectively (Savich and Samovol 1976). Mars 4 occulted the ionosphere during autumn season at solar zenith angle ~ 127°. The location was at 90° S, 236° W and the local time was 03:30 h. Mars 5 occulted the ionosphere during spring season at solar zenith angle ~ 106°. The location was at 38° N, 214° W and the local time was 04:30 h. Two electron density profiles were obtained from these measurements, which are shown in Fig. 20.1. Each profile shows the presence of two plasma layers, one at height ~ 80 km and other with a small shoulder of electron density ~ 1–2 × 103 cm−3 at 25 km (Molina-Cuberos et al. 2008; Haider et al. 2009). The ultraviolet spectrometer onboard Mars 5 found nothing in the upper atmosphere of Mars at wavelength range 0.3–0.8 μm. Only upper limits of

4.3 Viking 1 and 2

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(1) Optical electric devices of the echo navigation system, (2) Instrument module, (3) Container of altitude system, (4) Magnetometer, (5) Stereo antenna, (6) Highgain parabolic antenna, (7) Low- gain cone antenna, (8) Tanks of propulsion syste, (9) Radiators of temperature control system, (10) Nozzles of the altitude system, (11) Nozzles of stabilization system, (12) Solar panel and (13) Opticalelectronic device of positioning system Fig. 4.2 Mars 3 spacraft with scientific instruments flown at Mars from IKI, Russia (Credit IKI)

~ 50 Rayleigh were evaluated in the dark limb of the planet from Mars 5 spectroscopic observations (Karsnopolsky and Krysko 1975) (Fig. 4.2)

4.3 Viking 1 and 2 The Viking 1/2 spacecrafts carried the orbiters and Landers together. The orbiter carried the Atmospheric Water Detector (MAWD) instrument (Farmer et al. 1976), which is shown in Fig. 4.3a. The Lander carried the six aeronomy instruments: (1) Neutral Mass Spectrometer (NMS) (Nier and McElroy 1976), (2) Retarding Potential Analyser (RPA) (Hanson et al. 1977), (3) X-ray Fluroscence (XRF) spectrometer (Toulmin et al 1976) (4) Gas Chromatograph Mass Spectrometer (GCMS) (Rushneck et al. 1978), (5) Seismometer (Anderson et al. 1977) and (6) Radio Occultation (RO) experiment (Fjeldbo et al. 1977) (see Fig. 4.3 b). The antennas and other necessary components of the Viking spacecraft are also shown in these figures. The first direct measurements of the neutral atmosphere of Mars were accomplished by the two Viking Landers. Viking 1 landed on Mars on 20 July, 1976 at 22.5° N, 48° W about 4 h after local noon with solar zenith angle 44°. Viking 2 landed on Mars at 48° N, 22° W on 3 September, 1976 about 10 h after local

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Fig. 4.3 Image of Viking 1 (a) orbiter and Lander (b) shows the aeronomy instruments flown at Mars (Credit NASA)

4.4 Mars Global Surveyor (MGS)

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mid-night with solar zenith angle 44°. The NMS onboard Viking 1 and 2 measured the neutral compositions (CO2 , N2 , CO, O2 and NO) during their landings. The density of O could not be measured from this spectrometer. However, it was inferred from O2 + and CO2 + densities measured by RPA experiment onboard Viking 1 and 2. Helium has also been detected on Mars by Krasnopolsky et al. (1994) through airglow measurements. The RO measurements were also carried out from Viking 1 and 2 at solar zenith angle ranging from 45° to 127°. The sub-solar and mid-night ionosphere regions at Mars have never been probed. Like the plasmapause in Earth’s ionosphere a similar feature is observed in the Martian ionosphere. This feature has been given the name ionopause. Brace et al. (1983) defined it as the altitude where electron density abruptly drops to 102 cm−3 . In most of the radio occultation measurements of Mariners and Vikings, there is no abrupt drop in the electron density and it gradually approaches the measurements noise level. In these cases, the highest point is chosen as ionopause height. It should be noted that high ionopause heights often occur in the Viking data during solar maximum and intermediate conditions for solar zenith angles greater than 55° (Ness et al. 2000).

4.4 Mars Global Surveyor (MGS) MGS was launched on 7 November, 1996, which operated on Mars until 21 November, 2006. It carried five instruments: (1) Mars Orbiter Camera (MOC) (Malin and Edgett 2001), (2) Thermal Emission Spectrometer (TES) (Conrath et al. 2000), (3) Mars Orbiter Laser Altimeter (MOLA) (Garvin et al. 1999), (4) Magnetometer and Electron Reflectometer (MAG/ER) (Acuña et al. 1998) and (5) RO experiment (Albee et al. 2001). In addition, Mars horizon sensor and accelerometer instrument have been used to study the atmospheric dynamics of Mars. MGS provided large data sets of mass densities, magnetic fields and electron densities at various locations in the upper atmosphere of Mars (Bougher et al. 2003; Krymskii et al. 2003; Bougher et al. 2004 and Haider et al. 2006, 2010). It observed 5600 electron density profiles from RO experiment during its entire period (Withers et al. 2008, 2015; Fellows et al. 2015; Thirupathaiah et al. 2019; Shah et al. 2021).The mass density was measured by accelerometer experiment during the aerobraking period of MGS (Keating et al. 1998; Haider et al. 2010; Haider and Mahajan 2014). Aerobraking took place in two phases—Phase 1 and 2. Phase 1 included orbits #P1-P201, from mid September 1997 to late March 1998, while Phase 2 included the orbits #574-P1283 from mid September 1998 to early February, 1999. The magnetic field in the upper atmosphere of Mars was measured by the magnetometer onboard MGS (Connerney et al. 1999). The magnetic field as high as 400 nT at 108–113 km altitude in northern hemisphere and 1500 nT at 120–200 nT altitude in some locations of the southern hemisphere were observed from this spacecraft. However, the strength of the magnetic fields in other locations at the same altitude range is much lower about 5–15 nT. These magnetic fields are so weak that they do not

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significantly contribute to the solar wind—Mars interaction. Thus, solar wind interaction with Mars outside the crustal magnetic fields should be induced like Venus. As a consequence solar wind daytime pressure nearly permanently compresses the interplanetary magnetic field (IMF) into the Martian ionosphere down to 180 km by several orders of magnitude (Shinagawa and Cravens 1989; Shinagawa and Bougher 1999). Figure 4.4 represents the image of MGS spacecraft with five instruments flown at Mars. The spacecraft sensors, solar panel, and antennas are also shown in this figure.

4.5 Mars Express (MEX) The MEX was launched on 2 June 2003, which is operating till today. It carried seven instruments: (1) High Resolution Stereo Camera (HRSC) (Neukum and Jaumann 2004), (2) Analyser of Space Plasmas and Energetic Atoms (ASPERA) (Barabash et al. 2004), (3) Mars Radio Science (MaRs) (Pätzold et al. 2004) (4) Mars Advanced Radar for Subsurface and Ionosphere Sounding (MARSIS) (Gurnett et al. 2005), (5) Observatoire Pour La Mineralogie I’ Eau Las Glaces et’Active (OMEGA) (Bibring et al. 2004), (6) Planetary Fourier Spectrometer (PFS) (Formisano et al. 2005), and (7) Spectroscopy for Investigation of Characteristics of the Atmosphere of Mars (SPICAM) (Bertaux et al. 2006) for atmospheric studies. The MaRs instrument was carried out mostly by all aeronomy missions for the study of the ionosphere of Mars. The M1 and M2 peaks have been observed in the daytime ionosphere at altitude 110 and 135 km respectively from RO experiment onboard Mariner 6, 7 and 9 (Fjeldbo et al. 1970; Kliore et al. 1972), Mars 2, 3, 4 and 5 (Rasool and Stewart 1971) and Viking 1/2 (Fjeldbo et al. 1977). MGS and MEX have carried out further observations of three ionization peaks in most of the daytime electron density profiles at altitude ~ 130–140 km, 100–115 km and ~ 75– 90 km (Pätzold et al. 2005; Withers et al. 2008; Pandya and Haider 2014). Figure 4.5 represents three ionization peaks as M2, M1 and meteoric layers at Mars. The solar wind interacts directly above 200 km, where collision mean free path is large. The ionopause structure is also shown in Fig. 4.5 at about 350–400 km. MGS observed 5600 electron density profiles in the daytime ionosphere of Mars during the period 24 December 1998 to 9 June 2005 while MEX has reported 500 electron density profiles (Withers et al. 2012). Figure 4.6 represents image of MEX spacecraft carrying seven scientific instruments at Mars.

4.6 Mars Atmosphere and Volatile Evolution (MAVEN) MAVEN was launched on November 18, 2013 and it entered into the orbit of Mars on September 21, 2014. The mission goal is to explore the planet’s upper atmosphere, ionosphere and interaction with the sun and solar wind. The mission data is

Fig. 4.4 Image of MGS spacecraft with five instruments flown at Mars (Credit NASA)

4.6 Mars Atmosphere and Volatile Evolution (MAVEN) 25

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Fig. 4.5 Schematic representation of ionospheric layers at altitude < 200 km and above where the topside ionospheric structure are shown by wavy lines (from Withers 2009)

Fig. 4.6 Image of MEX spacecraft carrying seven scientific instruments at Mars (Credit NASA)

4.7 Mangalyaan

27

being used to study the loss of volatile compounds such as CO2 , N2 and H2 O from the Martian atmosphere to space. It carried nine instruments: (1) Neutral Gas and Ion Mass Spectrometer (NGIMS) (Mahaffy et al. 2015a, b), (2) Imaging Ultraviolet Spectrometer (IUVS) (McClintock et al. 2015), (3) Magnetometer (Connerney et al. 2015a, b), (4) Solar Wind Electron Analyzer (SWEA) (Mitchell et al. 2016), (5) Langmuir Probe and Waves Antena (LPWA) (Anderson et al. 2015), (6) Solar Energetic Particle (SEP) (Rahmati et al. 2015), (7) Solar Wind Ion Analyzer (SWIA) (Halekas et al. 2015a, b) and (8) Supra Thermal and Thermal Ion Composition (STATIC) (McFadden et al. 2015), and (9) Extreme Ultraviolet (EUV) monitor (Eparvier et al. 2015). Radio Occultation Science Experiment (ROSE) data is basically the signal transmitted from MAVEN to the ground so the instrument consists of the MAVEN radio transmitter and HGA and the antennas of the Deep Space Network (DSN). The NGIMS measures the composition and isotopes of thermal neutrals and ions. The IUVS measures global characteristics of the upper atmosphere and ionosphere via remote sensing observations (Schneider et al. 2015a, b). The magnetometer measures the planet’s magnetic field components Bx, By, and Bz in X, Y, and Z directions via flux gate sensors (Connerney et al. 2015a, b). Unlike the Earth’s magnetic field, which surrounds the entire planet, Mars only has patches of magnetic field left in the crust (Schneider et al. 2021). The SWEA is a symmetric hemispheric electrostatic analyzer, which is designed to measure the energy and angular distributions of solar wind electrons and ionospheric photoelectrons in the Mars’ environment (Mitchell et al. 2016). The LPWA instrument is designed to measure the electron density, temperature and electric field in the ionosphere of Mars (Andersson et al. 2015). The SEP instrument is a part of the particles and fields package, which determines the impact of SEPs on the upper atmosphere of Mars (Jakosky et al. 2015). The SWIA instrument measures the solar wind and magnetosheath proton flow around Mars and constrains the nature of solar wind interactions with the upper atmosphere (Halekas et al. 2017). The STATIC instrument is an electrostatic analyzer, which measures 3-D ion distributions and detects pick up ions at high energy in the Mars’ environment (Rahmati et al. 2018). The EUV monitor is a part of the LPWA instrument and measures the solar EUV variations and wave heating of the Martian upper atmosphere. Figure 4.7 shows the image of MAVEN spacecraft carrying nine scientific instruments at Mars.

4.7 Mangalyaan Mangalyaan was launched on November 5, 2013, which entered into the orbit of Mars on September 21, 2014. It has successfully completed the planned mission lifetime and is now extended mission phase. This spacecraft carried five instruments: (1) Lyman Alpha Photometer (LAP) (Sridhar Raja et al. 2015), (2) Mars Exospheric Neutral Composition Analyzer (MENCA) (Bhardwaj et al. 2015), Methane Sensor for Mars (MSM) (Mathew et al. 2015), (4) Mars Color Camera (MCC) (Arya et al. 2015) and (5) Thermal Infrared Imaging Spectrometer (TIS) (Singh et al. 2015).

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Fig. 4.7 Image of MAVEN spacecraft carrying nine scientific instruments at Mars (Credit NASA)

The LAP measures the relative abundance of deuterium and hydrogen (D/H ratio) based on absorption of Lyman α line of D and H in the upper atmosphere of Mars. The MENCA is a quadrapole mass spectrometer covering the mass range of 1–300 amu with mass resolution of 0.5 amu. The MSM is a Fabry–Perot Etalon sensor to measure CH4 at ppb (parts per billion) level and map its source (Mathew et al. 2015). The MCC is imaging topography of Mars at square area ~ 50 km × 50 km (Arya et al. 2015). The TIS is an infrared mapping spectrometer, which provides map of composition and mineralogy of the Martian surface (Singh et al. 2015). Figure 4.8 shows the image of Mangalyaan spacecraft carrying the five scientific instruments at Mars (Credit ISRO).

4.8 Emirates Mars Mission (EMM) The EMM was launched on 19 July, 2020 and it went into the orbit around Mars on 9 February, 2021. This mission is studying the Martian atmosphere and climate. It will also help to answer key questions about the global Martian atmosphere and the loss of hydrogen and oxygen gases into space. The EMM carried three scientific instruments: (1) Emirates Mars Ultraviolet Spectrograph (EMUS) (Holsclaw et al.

4.9 ExoMars Mission

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Fig. 4.8 Image of Mangalyaan spacecraft carrying five scientific instruments at Mars (Credit ISRO)

2021), (2) Emirates Exploration Imager (EXI) (Jones et al. 2021) and (3) Emirates Infrared Spectrometer (EMIRS) (Edwards et al. 2021). The EXI camera’s high resolution image is measuring the properties of Mars’ water ice dust, aerosols and ozone in its atmosphere. The EMUS measures the global characteristics and variability of hydrogen and oxygen in the Mars thermosphere. The EMUS instrument frame is oriented such that the + Z direction is parallel to the spacecraft + Y direction, while the instrument + Y direction is parallel to the spacecraft + X direction. The EMIRS is observing the distribution of dust particles and ice clouds and is also tracking the movement of water of water vapor and heat through the atmosphere. Figure 4.9 shows the image of EMM spacecraft carrying three scientific instruments at Mars (Amiri et al. 2022).

4.9 ExoMars Mission The goals of ExoMars are: (1) To search for signs of past life on Mars, (2) How the Martian water and geochemical environment varies, and (3) To investigate atmospheric trace gases and their sources. It will after so demonstrate the technologies for a future Mars Sample Return mission. The ExoMars was launched in 2016 that placed a TGO into the Mars orbit. It carried four instruments: (1) Nadir and Occultation for Mars Discovery (NOMAD) Vandaele et al. 2018a, b), (2) Atmospheric Chemistry Suite (ACS) (Korablev et al. 2018), (3) Color and Stereo Surface Imaging System (CaSSIS) (Tao et al. 2021), and (4) Fine Resolution Epithermal Neutron Detector (FREND) (Mitrofanov et al. 2018).

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Fig. 4.9 Image of EMM carrying three scientific instruments at Mars (from Amiri et al. 2022)

NOMAD comprises three spectrometers two sensitive for infrared and one for ultraviolet wave to identify atmospheric components such as methane and hydrocarbons via solar occultation measurements 2nd by studying radiation reflected by the Martian atmosphere (Vandaele et al. 2018a, b). ACS includes three infrared instruments to look at the chemistry and structure of the Martian atmosphere (Korablev et al. 2018). CaSSIS is delivering images of the Martian surface to be used in conjunction with atmospheric data to identify sources and sinks for trace gases such as methane, which is known to be present in certain geological locations, while not being detected in other areas (Thomas et al. 2017). This instrument is comprised of a high resolution camera reaching a ground resolution of 5 m per pixel for color and stereo imaging across a wide ground swarh. FREND is observing subsurface hydrogen to a depth of one meter to uncover water ice deposits near the surface (Semkova et al. 2018). The measurements of subsurface water and OH by FREND will be ten times better than previous missions. Recently, NOMAD instrument onboard ExoMars/TGO observed mixing ratios of H2 O and HDO in the lower atmosphere of Mars in the presence and absence of Global Dust Storm (GDS) 2018 (Vandaele et al. 2018a, b). The D2 O is not measured in the Martian atmosphere because its mixing ratio was below the detection limit of the instrument. The image of EXOMars TGO is shown in Fig. 4.10.

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Fig. 4.10 Image of EXOMars TGO mission carrying four scientific instruments at Mars (Credit ESA)

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Chapter 5

Thermal Structure of Mars Atmosphere

Abstract The atmosphere of Mars consists of troposphere, stratosphere, mesosphere, thermosphere and exosphere. Temperature profile of Mars depends on its orbital parameters and its atmospheric compositions. Mars’ average distance from the sun is 1.52 AU and its higher orbital eccentricity (0.093) causes significant changes in total received solar radiation. Figure 5.1 shows vertical profiles of the neutral temperature observed by Viking 1, 2 Landers, Curiosity, and Opportunity rovers. The mean surface temperature is about 220 K and it varies with time of the day, season and location on the planet from a minimum near 147 K in polar nights to a maximum around 300 K near the sub-solar point at perihelion (Barth et al., Annu. Rev. Earth Planet Sci. 2:333, 1974). Keywords Atmosphere · Temperature structure · Ionosphere

5.1 Troposphere The troposphere occurs between altitude 0 and 50 km. In this region the temperature is decreasing lapse rate is 1.6–1.8 K/km. The adiabatic lapse rate for Mars is ~ 5 K/ km (Haberle et al. 2017; Read and Lewis 2004), whereas, for Earth is ~ 9.8 K/km, which can be accounted by the thin atmosphere of Mars. In the troposphere most of the weather phenomena like convection and dust storms take place. The dynamics of the temperature is heavily driven by the daytime surface heating and the amount of suspended dust. Near-surface, diurnal temperature variability is large due to low thermal inertia. Under dusty condition, the suspended dust particles can reduce the surface diurnal temperature. The temperature above 15 km is controlled by radiative processes instead of convection (Fig. 5.1).

© Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_5

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Fig. 5.1 The vertical structure of the atmosphere of Mars overlaying with temperature profiles retrieved from the entry probes of Mars Landers (https://nasa.pds.ppi)

5.2 Stratosphere Mars does not have a persistent stratosphere due to the lack of shortwave-absorbing species in its middle atmosphere. However, a seasonal ozone layer and a strong temperature inversion in the middle atmosphere have been observed over the Martian

5.5 Exosphere

39

South pole (McCleese et al. 2008). The altitude of the turbopause of Mars varies greatly from 60 to 140 km, and the variability is driven by the CO2 density in the lower thermosphere.

5.3 Mesosphere Mars has a mesosphere (~ 50 to ~ 100 km) just above the troposphere, which is defined on the basis of vertical temperature structure. The lapse rate is lower in the mesosphere than the troposphere. In this region the temperature is almost constant because of insufficient O3 to absorb UV radiation. Temperature in the mesosphere is greatly under the influence of tides and waves. The CO2 in the mesosphere acts as a cooling agent by efficiently radiating heat into space. Stellar occultation observations show that the mesopause of Mars is located at about 100 km (around 0.01–0.001 Pa level) and has a temperature ~ 100–120 K.

5.4 Thermosphere There is a thermosphere between 100 and 200 km, where temperature increases with altitude due to absorption of solar EUV and X-ray radiation. The temperature of the Martian thermosphere increases with altitude and varies with seasons. The daytime temperature of the upper thermosphere of Mars varies from 175 K (at aphelion) to 240 K (at perihelion), which is significantly lower than the temperature of Earth’s thermosphere.

5.5 Exosphere Above ~ 200 km, the atmospheric density is almost negligible, leading to the insignificant change in the temperature. This region is known as exosphere (~200 km and above). The temperature in the exosphere is nearly constant and atmospheric species escape from the exobase. Thus collision is almost negligible in the exosphere of Mars. The exobase is defined at that altitude where the neutral scale height is equal to the mean free path. Because of the smaller scale height in the Martian thermosphere, exobase occurs at a smaller altitude in comparison to the exobase of Earth’s atmosphere. In the lower atmosphere H2 is produced due to photo dissociation of H2 O and diffuses to the exosphere. The exospheric H2 then decomposes into hydrogen atom that have sufficient thermal energy can escape from the gravitation of Mars. The escape of atomic hydrogen is evident from the UV spectrometer on different orbiters (Chaffin et al. 2015). Recently hydrogen exosphere of Mars is observed from MAVEN

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(Halekas 2017). The escape of atmospheric gases from Mars plays a key role in transforming Mars from a warm and wet early state to its current cold dry conditions.

References Barth, C.A.: The atmosphere of Mars. Annu. Rev. Earth Planet Sci. 2, 333 (1974) Chaffin, M.S., Chaufray, J.Y., Deighan, J., et al.: Three-dimensional structure in the Mars H corona revealed by IUVS on MAVEN. Geophys. Res. Lett. 42(21), 9001–9008 (2015) Haberle, R.M., Clancy, R.T., Forget, F., et al.: The Atmosphere and Climate of Mars. Cambridge University Press (2017) Halekas, J.S.: Seasonal variability of the hydrogen exosphere of Mars. J. Geophys. Res. Planets 122(5), 901–911 (2017) McCleese, D.J., Schofield, J.T., Taylor, F.W., et al.: Intense polar temperature inversion in the middle atmosphere on Mars. Nat. Geosci. 1(11), 745–749 (2008) Read, P.L., Lewis, S.R.: The Martian Climate Revisited: Atmosphere and Environment of a Desert Planet. Springer Science and Business Media (2004)

Chapter 6

Magnetic Field of Mars

Abstract The first observations of magnetic fields and plasmas in the near Mars environment were obtained by Mariner 4 in 1998 (Anderson Science 149:1226– 1228, 1965; Acuña et al. Science 279:1676–1680, 1998). Later observations were carried out by Mars 2 to 7 and Phobos − 2 in 1970 and 1989 respectively (Riedler et al. Planet Space Sci. 39:75–81, 1991; Slavin et al. J. Geophys. Res. Space Phys. 96:11235–11241, 1991; Möhlmann et al. Adv Space Res 12:221–229, 1992). These measurements were carried out above the atmosphere/ionosphere of Mars. After these missions, MGS was the first spacecraft to obtain the magnetic field observations beneath the ionosphere (~ 170 to 200 km). MEX, EMM and ExoMars did not carry its own magnetometer on Mars. MAVEN also carried a magnetometer, which measures the magnetic fields in the upper atmosphere of Mars. This instrument helps us to see where the atmosphere is protected by mini-magnetosphere and where it is open to solar wind. Keywords Magnetosphere · Induced · Intrinsic magnetic fields

6.1 Earlier Measurements: Mariner 4, Mars 2 to 7 and Phobos 2 Mariner 4 and Mars 2 to 7 have observed upper limits on the Martian magnetic moment in the range of 1–5 × 1012 Tm3 corresponding to an equatorial surface field of 25–125 nT (Dolginov et al. 1973, 1978a, 1978b, 1978; Riedler et al. 1989). The average distance of Martian bow shock is ~ 2.66 ± 0.05 Rm (where Rm is the radius of Mars ~ 3390 km). Viking Landers reached on the surface of Mars in 1976 but did not carry the magnetic field experiment (Snyder and Moroz 1992). Phobos 2 also observed the bow shock at Mars (Slavin et al. 1991). The Martian bow shock location is independent of solar cycle phase, hence, solar cycle EUV flux. The magnetic fields observed by Phobos 2 have suggested that the intrinsic field is very weak at the bow shock and Mars should have a hybrid magnetosphere (combination of intrinsic and induced magnetic fields) (Vaisberg and Smirnov 1986). The location of bow shock

© Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_6

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has been studied extensively from these missions. Their observations confirmed subsolar (nose) radii at ~ 1.36–1.85 Rm and terminator radii at ~ 2.36–3.07 Rm (Riedler et al. 1989).

6.2 Latest Measurements: MGS and MAVEN The magnetometer onboard MGS observed the magnetic field as high as 400 nT at 108–113 km altitude in northern hemisphere and 1500 nT at 120–200 km altitude in some locations of southern hemisphere from this spacecraft (Acuña et al. 1998; Connerney et al. 1999, 2005). However, the strength of the magnetic fields in other locations at the same altitude range is much lower about 5–15 nT. These magnetic fields are so weak that do not significantly contribute to the solar wind—Mars interaction. Thus, the solar wind interaction with Mars outside the crustal magnetic field region should be induced like Venus. As a consequence, solar wind dynamic pressure nearly permanently compresses the IMF into the Martian ionosphere down to 180 km by several orders of magnitude (Shinagawa and Cravens 1989). MAVEN magnetometer also measures the magnetic fields of Mars above 150 km (Connerney et al. 2015a, b . In Fig. 6.1 five years data of this magnetometer has been used to create a global map of the electric current system in the induced magnetosphere (Ramstad et al. 2020). These currents play an important role in the atmospheric loss from Mars. The Earth has also such current system, which can be seen in the form of colorful displays of light in the night sky near polar region known as aurora. The Earth’s aurora are strongly linked to such currents generated by interaction of Earth’s magnetic field with solar wind that flow along vertical magnetic field lines into the atmosphere concentrating in the polar regions. Since Mars is a rocky terrestrial planet, one might assume that same kind of magnetic paradigm functions can also occur there. However, Mars does not generate a magnetic field on its own, outside of relatively small patches of magnetized crust. Something different from what we observe on Earth must be happening on Mars. In Fig. 6.1 Generator currents (E.J < 0) are colored blue and load currents (E.J > 0) are colored red. The magnetic pressure around the planet induces Chapman-Ferraro type currents at the induced magnetosphere boundary (JIMB ) as well as bow shock currents (JBS ). The southward flow of charge sets up a potential difference (φ+ − φ− ) across the system and a resulting magnetospheric convective electric field, in turn, drives currents in the Ohmic conductive ionosphere (JIono ) and the gyro conductive tail (JTail ). A twist is present in the dayside ionospheric and tail currents, related to Sunward/anti-Sunward component present along the southern sheath the IMB flanks the terminator ionosphere and the inner tail. VSW is the solar wind velocity vector, BIMF is the upstream IMF field vector, and Emot is the motional electric field of the solar wind. The induced magnetosphere is represented by Mars-Sun-Electric field (MSE) coordinates (Dubinin et al. 1996). In MSE coordinates XMSE is anti-parallel to the solar wind flow vector (VSW ), ZMSE defines the orientation of the solar wind

6.3 Induced and Crustal Magnetic Fields

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Fig. 6.1 Illustration of formative current systems in the Martian induced magnetosphere (from Ramstad et al. 2020)

motional electric field, and YMSE completes this right-handed system in the direction of the IMF component orthogonal to VSW and Emot .

6.3 Induced and Crustal Magnetic Fields Brain et al. (2003) has reported that weak crustal fields extend above 120 km altitude over ~ 70% of the surface. The strongest crustal magnetic fields extend above 1000 km altitude over some locations of the southern surface (e.g., 140° E to 240° E) (Haider et al. 2010). The strengths of the induced magnetic fields are about 5–15 nT in other locations outside the crustal magnetic field region. The induced magnetic fields are so weak that they do not significantly contribute to the solar wind-Mars interaction. These magnetic fields have two components: (1) horizontal component, BH and (2) radial component, BR . Total magnetic field, B is equal to the square root of the sum of

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Fig. 6.2 Altitude profiles of magnetic field components (BR , BH and elevation angle) in the dayside ionosphere of Mars over Olympus Mons. Cain’s model predictions are also shown by red dot lines (from Haider et al. 2010)

B2 H and B2 R . The elevation angle of the magnetic field with respect to local horizontal is obtained as tan θ = BR /BH . In Figs. 6.2 and 6.3 the altitude profiles of BR , BH and elevation angles are shown for Olympus Mons (15° N, 210° E) and Syrtis Major (5° N, 30° E), respectively. These profiles were measured from MGS magnetometer in the dayside atmosphere of Mars (Haider et al. 2010). The measurements are compared with the crustal magnetic field model of Cain et al. (2003). The maximum value of horizontal magnetic field of ~ 125 nT is measured at 112 km over Olympus Mons. Similarly the highest value of vertical magnetic field of 400 nT is observed at ~ 112 km over Syrtis Major. The crustal magnetic field strength decreases rapidly in the Martian atmosphere above 112 km.

6.4 Mini-Magnetosphere At Mars, the conductive ionosphere and localized crustal magnetic anomalies interact with the upstream solar wind creating a unique, hybrid magnetosphere (Bowers et al. 2021). As the IMF lines encounter Mars, the nature of their interaction with the magnetosphere will vary depending on whether they encounter regions by crustal anomalies are absent, which occur primarily in northern hemisphere or regions of strong (100–1000 nT at the surface) crustal magnetic fields, which primarily in the southern hemisphere. Around regions of weak crustal magnetic fields, the IMF first encounters the dayside ionosphere of Mars and drapes around the obstacle in absence of global intrinsic magnetosphere. Around regions of strong

6.5 Global Modeling of Mars Magnetic Field

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Fig. 6.3 Altitude profiles of magnetic field components (BR , BH and elevation angle) in the dayside ionosphere of Mars over Syrtis Major. Cain’s model predictions are also shown by red dot lines (from Haider et al. 2010)

crustal fields, the magnetic anomalies protrude out into space and create a “minimagnetospheric” interaction between the draped IMF and the crustal anomalies analogous to the dayside of Mars with a global intrinsic dipole field. Previous studies at Mars using both the MGS and MAVEN observations have found evidence of complex dynamics within these mini-magnetospheres during the dayside ionosphere, including current sheets, and magnetic flux ropes. Both these phenomena play a role in plasma transport and acceleration within the ionosphere and are associated with magnetic reconnection, which implicates a variety of dynamic processes in Martian mini-magnetospheres.

6.5 Global Modeling of Mars Magnetic Field Recently Martian magnetic field of Mars is estimated by Gao et al. (2021) using a Spherical Harmonic (SH) model. In this model MGS and MAVEN datasets have been used in Planetocentric Coordinate (PC) system. Longitude, latitude and spherical coordinates (r, θ, φ) are defined within the frame of PC, where r is radial distance, θ points southward and φ points eastwrd. The magnetic field data in PC Cartesian coordinate are transformed to PC spherical coordinate, which have three components Br (radial), Bθ (southward), and Bφ (eastward). The radius of Mars is set to 3393.5 km. In SH model the magnetic potential satisfies Laplace’s equation ∇ 2 V = 0, where V is a magnetic potential. The V has two components describing internal crustal field and external induced field potentials Vi and Ve , respectively. Vi can be expressed as follows (Gao et al. 2021):

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Vi (r, θ, φ) = a

n N ( ) ∑ a n+1 ∑ ( n=1

r

) m gnm cos mφ + h m n sin mφ Pn (cos θ )

(6.1)

m=0

B→ = −∇V

(6.2)

N ( ) n ∑ ∑ ( m ) m a n+2 −∂ V gn cos mφ + h m Br = (n + 1) = n sin mφ Pn (cos θ ) ∂r r n=1 m=0

(6.3) Bθ =

n N ( ) ∑ ) ∂ Pnm (cos θ ) a n+2 ∑ ( m −∂ V gn cos mφ + h m =− n sin mφ r ∂θ r ∂θ n=1 m=0

−∂ V r sin θ ∂φ n N ) m 1 ∑ ( a )n+2 ∑ ( m −gn sin mφ + h m =− n cos mφ Pn (cos θ ) sin θ n=1 r m=0 √ B = Br2 + Bθ2 + Bφ2

(6.4)

Bφ =

(6.5)

(6.6)

where N is maximum SH degree and a represents radius of Mars, r, θ, and φ are the radial distance, latitude and longitude of a given location, respectively. Pnm is the Schmidt semi-normalized associated Legendre function for nth degree and mth order, gnm and h m n are the SH coefficients which are also called Gauss coefficients. Using SH model magnetic field B, and its three components are derived from Eqs. (6.3) to (6.6). The larger degree N represents the better resolution of the model (Gao et al. 2021). Using this model Gao et al. (2021) calculated crustal magnetic fields at an altitude ~ 120 km. The calculated magnetic fields of Br , Bθ , and Bφ are shown in the map of Fig. 6.4, where the black lines represents the zero altitude of the Mars Orbiter Laser Altimeter (MOLA) topography (Smith et al. 2001). This model can be used to predict the crustal fields at any given altitude above the Martian surface (Gao et al. 2021). There are several regions in the northern hemisphere, which are devoid of intense crustal magnetic fields. The most intense crustal fields are estimated as 200–400 nT in southern hemisphere at near east longitude ~ 180o E.

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Fig. 6.4 Global crustal magnetic field maps of Mars at an altitude of 120 km (from Gao et al. 2021)

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References Acuña, M.H., et al.: Magnetic field and plasma observations at Mars: initial results of the Mars global surveyor mission. Science 279(5357), 1676–1680 (1998) Anderson, H.R.: Mariner IV measurements near mars: initial results. Science 149(3689), 1226–1228 (1965) Bowers, C., Slavin, J., DiBracccio, G., et al.: MAVEN observations of mini-magnetosphere dynamics within the dayside ionosphere of Mars. (No EPSC2021-305) Copernicus Meetings (2021) Brain, D.A., Bagenal, F., Acuña, M.H., Connerney, J.E.P.: Martian magnetic morphology: contributions from the solar wind and crust. J. Geophys. Res. Space Phys. 108(A12) (2003) Cain, J.C., Ferguson, B.B., Mozzoni, D.: An n= 90 internal potential function of the Martian crustal magnetic field. J. Geophys. Res. Planets 108(E2) (2003) Connerney, J.E.P., Acuna, M.H., Wasilewski, P.J.P.A., et al.: Magnetic lineations in the ancient crust of Mars. Science 284(5415), 794–798 (1999) Connerney, J.E.P., Acuña, M.H., Ness, N.F., et al.: Tectonic implications of Mars crustal magnetism. Proc. Natl. Acad. Sci. 102(42), 14970–14975 (2005) Connerney, J.E.P., Espley, J., Lawton, P., et al.: The MAVEN magnetic field investigation. Space Sci. Rev. 195(1), 257–291 (2015a) Connerney, J.E., Espley, J.R., DiBraccio, G.A., et al.: First results of the MAVEN magnetic field investigation. Geophys. Res. Lett. 42(21), 8819–8827 (2015b) Dolginov, S.S.: On the magnetic field of Mars: Mars 2 and 3 evidence. Geophys. Res. Lett. 5(1), 89–92 (1978b) Dolginov, S.S.: On the magnetic field of Mars: Mars 5 evidence. Geophys. Res. Lett. 5(1), 93–95 (1978c) Dolginov, S.S., Yeroshenko, Y.G., Zhuzgov, L.N.: Magnetic field in the very close neighborhood of Mars according to data from the Mars 2 and Mars 3 spacecraft. J. Geophys. Res. Space 78(22), 4779–4786 (1973) Dolginov, S.: Magnetic fields in the vicinity of Venus according to” Venera” and” Mariner” data. In Origins of Planetary Magnetism, vol. 348, p. 22 (1978a) Dubinin, E., Sauer, K., Lundin, R., et al.: Plasma characteristics of the boundary layer in the Martian magnetosphere. J. Geophys. Res. Space Phys. 101(A12), 27061–27075 (1996) Gao, J.W., Rong, Z.J., Klinge, L., et al.: A spherical harmonic Martian crustal magnetic field model combining data sets of MAVEN and MGS. Earth Space Sci. 8(10), e2021EA001860 (2021) Haider, S.A., Seth, S.P., Brain, D.A., et al.: Modeling photoelectron transport in the Martian ionosphere at Olympus Mons and Syrtis major: MGS observations. J. Geophys. Res. Space Phys. 115(A8) (2010) Möhlmann, D., Sauer, K., Roatsch, T., et al.: Magnetic field environment of Mars as studied by Phobos-2. Adv Space Res 12(9), 221–229 (1992) Ramstad, R., Brain, D.A., Dong, Y.: The global current systems of the Martian induced magnetosphere. Nat. Astron. 4(10), 979–985 (2020) Riedler, W., Schwingenschuh, K., Lichtenegger, H., et al.: Interaction of the solar wind with the planet Mars: Phobos 2 magnetic field observations. Planet Space Sci. 39(1–2), 75–81 (1991) Riedler, W., Möhlmann, D., Oraevsky, V.N., et al.: Magnetic fields near Mars: first results. Nature 341(6243), 604–607 (1989) Shinagawa, H., Cravens, T.E.: A one-dimensional multispecies magnetohydrodynamic model of the dayside ionosphere of Mars. J. Geophys. Res. Space Phys. 94(A6), 6506–6516 (1989) Slavin, J.A., Schwingenschuh, K., Riedler, W., Yeroshenko, Y.: The solar wind interaction with Mars: Mariner 4, Mars 2, Mars 3, Mars 5, and Phobos 2 observations of bow shock position and shape. J. Geophys. Res. Space Phys. 96(A7), 11235–11241 (1991) Smith, David E., et al.: Mars Orbiter Laser Altimeter: Experiment summary after the first year of global mapping of Mars. J. Geophys. Res. Planets 106(E10), 23689–23722 (2001) Snyder, C.W., Moroz, V.I.: Spacecraft exploration of Mars. Mars 71–119 (1992) Vaisberg, O., Smirnov, V.: The martian magnetotail. Adv. Space Res. 6(1), 301–314 (1986)

Chapter 7

Upper Atmosphere of Mars

Abstract Mars’ atmosphere above 100 km is designated as the upper atmosphere where molecular diffusion dominates (McElroy et al., Science 194:1295–1298, 1976). It is primarily composed of CO2 (95%), N2 (2.8%) and Ar (2%) (Villanueva et al., Icarus 223:11–27, 2013); (Franz et al., Planet Space Sci. 138:44–54, 2017). It also contains trace gases like H2 O, O, CO, H2 and noble gases (Villanueva et al., Icarus 223:11–27, 2013); (Franz et al., Planet Space Sci. 138:44–54, 2017). The atmosphere of Mars is much thinner than the Earth’s atmosphere. The average surface pressure is about 6.0 mb, which is less than 1% of Earth’s pressure. Currently, thin Martian atmosphere does not have liquid water on the surface but many studies suggest that the Martian atmosphere was much thicker in the past (Catling, D.C. Mars atmosphere: history and surface interactions. In Encyclopedia of the Solar System, pp. 343–357. Elsevier (2014)). The atmosphere of Mars has been losing mass to space since the planet formed and the leakage of gases still continues today (Jakosky et al., Icarus 315:146–157, 2018). Dust devils and dust storms are prevalent on Mars, which occur on average every 3 Martian Years (MY). The mechanism of dust storm is still not well understood. Keywords Upper atmosphere · Thermosphere · Observations

7.1 MAVEN Measurement: Neutral Compositions The NGIMS onboard MAVEN measures neutral composition profiles in the daytime as well as in the nighttime atmosphere of Mars (Mahaffy et al. 2015a, b; Elrod et al. 2017). In Fig. 7.1 we have plotted sample profiles of CO2 , Ar, N2 , CO, O, and He as observed by NGIMS on 10 and 14 September, 2017, when the spacecraft was crossing from inbound (left panel plot) to outbound (right panel plot) orbits. MAVEN did not measure the neutral densities at altitudes below 150 km except during nine Deep Dip periods (e.g. Bougher et al. 2015b; Zurek et al. 2017; Stone et al. 2018). The neutral densities are nearly the same during inbound and outbound crossings of Mars atmosphere. These densities are decreasing exponentially with height, except for Helium, as expected. © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_7

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Fig. 7.1 MAVEN observations of neutral densities in the upper atmosphere of Mars corresponding to inbound and outbound orbits on 10 and 14 September, 2017 (from Haider et al. 2023)

The neutral densities observed by MAVEN are nearly same with the earlier measurements made by Viking 1/2 Landers. Viking Landers observed neutral density profiles at one location. MAVEN observed neutral densities at different locations. Therefore, we can study the latitudinal/longitudinal variability, effects of solar flares and dust storm on the neutral densities in the upper atmosphere of Mars. Recently, Solar flares provided a quantifiable transient perturbation to Mars atmosphere from its equilibrium state (Thiemann et al. 2015, 2018), which can be used to better constrain atmospheric models and a detailed modelling work using MTGCM can be carried out later (Bougher et al. 2009, 2015a; b).

7.2 Atmospheric Waves and Dynamics The dynamics of the Martian atmosphere is characterized mainly by the behavior of CO2 , water vapor, and dust in response to solar and seasonal variabilities and their interaction with the surface (Barnes et al. 2017). The atmosphere of Mars is also affected due to tidal, planetary and gravity wave oscillations. Atmospheric tides are global scale periodic oscillations of the atmosphere. They are persistent global oscillations that are observed in all types of atmospheric fields including wind, temperature, pressure, density and geo-potential height. The diurnal and semidiurnal tides have 24 h and 12 h periods, respectively. The oscillations related to long planetary

7.3 Photochemical Equilibrium on Mars

51

waves and baroclinic instability processes have been studied in the upper atmosphere of Mars (Seth et al. 2006a, b; Seth and Brahmanand Rao 2008). Mars is a windy planet, therefore, various types of atmospheric motions are expected to exist in the troposphere of Mars. The wind is composed of GW, tidal waves, and stationary waves. At high altitudes (~100–120 km) the GW dissipate rapidly due to molecular viscosity and thermal conduction because of their short wavelengths (Imamura and Ogawa 1995). Larger scale winds such as field oscillations are thought to be a dominant mode of the motion in the thermosphere of Mars (Bougher et al. 2001; Mahajan et al. 2007; Haider et al. 2009). The characteristic of longitudinal distribution of temperature and density in the Martian troposphere is assumed to be approximated as follows (Haider et al. 2009): Z = Co + C1 Sinϕ + C1 Cosϕ + · · · Cn Sin(nφ) + Cn Cos(nφ)

(7.1)

where ϕ is east longitude and Z is a function of ϕ. The parameters C0 , C1 ,C2 ……Cn are calculated using this equation in the matrix form Z = A C, where A is coefficient matrix and C is column matrix of parameters C1 to Cn. This approach is used by Bougher et al. (2001); Mahajan et al. (2007); Haider et al. (2009). for a least squares spectral fitting in the calculations based on observational data to characterize the zonal structures in the Martiant roposphere (Haider et al. 2009) for studying longitudinal variations in the thermosphere/ionosphere of Mars. Their model contains a constant temperature or density term, an amplitude and phase for a sinusoid with one cycle per 360°of the longitude which is labeled as wave 1 harmonic and higher harmonics up to and including wave 3 and so on. For waves 2 and 3 it will have 5 and 7 parameters respectively, which are different for each curve. The wave features indicate the presence of nonmigrating tidal variations present in the Mars thermosphere (Bougher et al. 2001).

7.3 Photochemical Equilibrium on Mars In the upper atmosphere photochemical model can be studied in absence of molecular diffusion. The production and loss rate in the chemical model calculates the neutral densities in the upper atmosphere. This calculation is valid up to ~ 200 km. The densities of major atmospheric gases like CO2 , N2 , O2 , Ar and O can be estimated from photochemical model. To estimate the complete altitude profile of neutral densities 1-D and 3-D transport models are developed.

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7.4 Upper Atmospheric Modeling The structure of the Martian upper atmosphere is complicated and therefore cannot be described by 1-D thermosphere/ionosphere models. Bougher et al. (2009) and Valeille et al. (2009) have developed 3-D Mars Thermosphere Global Circulation Model (MTGCM) and 3-D Direct Simulation Monte Carlo (DSMC) models, respectively to study the variation of thermosphere and exosphere of Mars at different local times, latitudes and longitudes. A combination of these models describes selfconsistently the Martian upper thermosphere/ionosphere and exosphere globally (Valeille et al. 2009). The model solves a finite difference primitive equation that self-consistently calculates neutrals, ion and electron densities over the globe under solar minimum, moderate and maximum conditions for different Mars seasons. The prognostic equations for neutral species CO2 , N2 , O2 , O and CO are included in this model. The zonal velocity, vertical velocity, temperature and geo-potential heights are also obtained on 33 pressure levels (above 1.32 μbar) corresponding to altitudes from 70 to 300 km with a 5° resolution over longitude and latitude. The vertical coordinate is logpressure with a vertical spacing of two grid points per scale height. In MTGCM model the parameters like f10.7 index (solar X-ray/EUV/UV flux variation), heliocentric distance (orbital variation) and solar declination (seasonal variation) can be varied. This model is also modified to accommodate atmospheric inflation, semidiurnal, diurnal tidal modes and phases consistent with dusty conditions (Bougher et al. 2006). The 1-D thermosphere models include both chemistry and transport by eddy diffusion and do not assume a fixed homopause (e.g. Nagy et al. 1980; Krasnopolsky 2002; Fox and Yeager 2006; Fox 2009). In early 1-D models of Mars thermosphere a single homopause altitude was assumed (e.g. McElroy 1967; Kumar and Hunten 1974; Chen et al. 1978). Below the homopause the atmosphere was considered to be completely mixed for chemical species. Above the homopause the neutral densities were assumed to be distributed according to their own scale heights. The species that are formed photochemically do not exhibit this behavior. The current 1-D models of Mars thermosphere include photoionization/excitation, photoelectron impact ionization/excitation, photodissociative excitation/ionization and more than 200 chemical reactions (e.g., Fox 2009). Recently Liu et al. (2018) have reported that the density of CO2 increased in the upper atmosphere of Mars up to ~ 200% in response to dust increases in the lower atmosphere. These dust enhancements occurred during the period from 20 December 2016 to 30 May 2017 at Ls ~ 320°. This study is consistent with the earlier findings when the whole atmosphere of Mars expanded and rose up to ~ 160 km due to the influence of dust (Withers and Pratt 2013). In Fig. 7.2 we show CO2 density profile obtained from MAVEN during the thermospheric warming associated with the planet-wide dust storm of 2018 (Elrod et al. 2020). In this figure we have also plotted air density from MCD for cold and warm scenarios of Mars’ atmosphere,

References

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200 CO 2 (NGIMS/MAVEN) MCD (Warm Atmosphere) MCD (Cold Atmosphere)

180

Altitude (km)

160 140 120 100 80 60 10 8

10 9

1010

1011

1012

1013

1014

1015

Neutral density (cm-3 )

Fig. 7.2 Altitude profiles of air densities estimated from MCD in the cold and warm atmosphere of Mars. The density of CO2 during thermospheric warming observed by NGIMS/MAVEN is also plotted (from Liu et al. 2018)

calculated for the global dust storm of 2007, which occurred during MY 28. It can be noted that in the warm scenario the air density is larger than the cold scenario by a factor of ~ 2–3.

References Barnes J.R., Haberle R.M., Wilson R.J., Lewis S.R., Murphy J.R., Read P.L.: The Atmosphere and Climate of Mars, ed. by Haberle R.M., Clancy R.T., Forget F., Smith M.D., Zurek R.W. Cambridge University Press, Cambridge. (2017) Bougher, S.W., Engel, S., Hinson, D.P., Forbes, J.M.: Mars Global Surveyor radio science electron density profiles: neutral atmosphere implications. Geophys. Res. Lett. 28(16), 3091–3094 (2001) Bougher, S.W., Pawlowski, D., Bell, J.M., et al.: Mars global ionosphere-thermosphere model: solar cycle, seasonal, and diurnal variations of the mars upper atmosphere. J. Geophys. Res. Planets 120(2), 311–342 (2015a) Bougher, S.W., Bell, J.M., Murphy, J.R., et al.: Polar warming in the Mars thermosphere: seasonal variations owing to changing insolation and dust distributions. Geophys. Res. Lett. 33(2) (2006) Bougher, S.W., McDunn, T.M., Zoldak, K.A., Forbes, J.M.: Solar cycle variability of Mars dayside exospheric temperatures: model evaluation of underlying thermal balances. Geophys. Res. Lett. 36(5) (2009) Bougher, S., Jakosky, B., Halekas, J., et al.: Early MAVEN deep dip campaign reveals thermosphere and ionosphere variability. Science 350(6261), aad0459 (2015b) Catling, D.C.: Mars atmosphere: history and surface interactions. In Encyclopedia of the Solar System, pp. 343–357. Elsevier (2014) Chen, R.H., Cravens, T.E., Nagy, A.F.: The Martian ionosphere in light of the Viking observations. J. Geophys. Res. Space Phys. 83(A8), 3871–3876 (1978)

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Elrod, M.K., Bougher, S., Bell, J., et al.: He bulge revealed: He and CO2 diurnal and seasonal variations in the upper atmosphere of Mars as detected by MAVEN NGIMS. J. Geophys. Res. Space Phys. 122(2), 2564–2573 (2017) Elrod, M.K., Bougher, S.W., Roeten, K., et al.: Structural and compositional changes in the upper atmosphere related to the PEDE-2018 dust event on Mars as observed by MAVEN NGIMS. Geophys. Res. Lett. 47(4), e2019GL084378 (2020) Fox, J.L., Yeager, K.E.: Morphology of the near-terminator Martian ionosphere: a comparison of models and data. J. Geophys. Res. Space Phys. 111(A10) (2006) Fox, J.L.: Morphology of the dayside ionosphere of Mars: implications for ion outflows. J. Geophys. Res. Planets 114(E12) (2009) Franz, H.B., Trainer, M.G., Malespin, C.A., et al.: Initial SAM calibration gas experiments on Mars: Quadrupole mass spectrometer results and implications. Planet Space Sci. 138, 44–54 (2017) Haider, S.A., et al.: Zonal wave structures in the nighttime tropospheric density and temperature and in the D region ionosphere over Mars: Modeling and observations. J. Geophys. Res. Space Phys. 114(A12) (2009) Haider, S.A., Durga, K.P., Shah, S.Y.: The magnetically controlled ionopause boundary observed by LPW onboard MAVEN within magnetic pile-up region of Mars. Icarus 394, 115423 (2023) Imamura, T., Ogawa, T.: Radiative damping of gravity waves in the terrestrial planetary atmospheres. Geophys. Res. Lett. 22(3), 267–270 (1995) Jakosky, B.M., Brain, D., Chaffin, M., et al.: Loss of the Martian atmosphere to space: present-day loss rates determined from MAVEN observations and integrated loss through time. Icarus 315, 146–157 (2018) Krasnopolsky, V.A.: Mars’ upper atmosphere and ionosphere at low, medium, and high solar activities: Implications for evolution of water. J. Geophys. Res. Planets 107(E12), 11–21 (2002) Kumar, S., Hunten, D.M.: Venus: an ionospheric model with an exospheric temperature of 350 K. J. Geophys. Res. 79(16), 2529–2532 (1974) Liu, G., England, S.L., Lillis, R.J., et al.: Thermospheric expansion associated with dust increase in the lower atmosphere on Mars observed by MAVEN/NGIMS. Geophys. Res. Lett. 45(7), 2901–2910 (2018) Mahaffy, P.R., Benna, M., Elrod, M., et al.: Structure and composition of the neutral upper atmosphere of Mars from the MAVEN NGIMS investigation. Geophys Res Lett 42(21), 8951–8957 (2015a) Mahaffy, P.R., Benna, M., King, T., et al.: The neutral gas and ion mass spectrometer on the Mars atmosphere and volatile evolution mission. Space Sci. Rev. 195(1), 49–73 (2015b) Mahajan, K.K., Singh, S., Kumar, A., Raghuvanshi, S., Haider, S.A.: Mars Global Surveyor radio science electron density profiles: some anomalous features in the Martian ionosphere. J. Geophys. Res. Planets 112(E10) (2007) McElroy, M.B.: The upper atmosphere of Mars. Astrophys. J. 150, 1125 (1967) McElroy, M.B., Kong, T.Y., Yung, Y.L., Nier, A.O.: Composition and structure of the Martian upper atmosphere: analysis of results from Viking. Science 194(4271), 1295–1298 (1976) Nagy, A.F., Cravens, T.E., Smith, S.G., et al.: Model calculations of the dayside ionosphere of Venus: ionic composition. J. Geophys. Res. Space Phys. 85(A13), 7795–7801 (1980) Seth, S.P., Brahmananda Rao, V.: Evidence of baroclinic waves in the upper atmosphere of Mars using the Mars Global Surveyor accelerometer data. J. Geophys. Res. Space Phys. 113(A10) (2008) Seth, S.P., Jayanthi, U.B., Haider, S.A.: Estimation of peak electron density in upper ionosphere of Mars at high latitude (50–70 N) using MGS ACC data. Geophys. Res. Lett. 33(19) (2006a) Seth, S.P., Brahmananda Rao, V., Esprito Santo, C.M., et al.: Zonal variations of peak ionization rates in upper atmosphere of Mars at high latitude using Mars Global Surveyor accelerometer data. J. Geophys. Res. Space Phys. 111(A9) (2006b) Stone, S.W., Yelle, R.V., Benna, M., et al.: Thermal structure of the Martian upper atmosphere from MAVEN NGIMS. J. Geophys. Res. Planets 123(11), 2842–2867 (2018)

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Thiemann, E.M.B., Eparvier, F.G., Andersson, L.A., et al.: Neutral density response to solar flares at Mars. Geophys. Res. Lett. 42(21), 8986–8992 (2015) Thiemann, E.M.B., Andersson, L., Lillis, R., et al.: The Mars topside ionosphere response to the X8 2 solar flare of 10 September 2017. Geophys. Res. Lett. 45(16), 8005–8013 (2018) Valeille, A., Combi, M.R., Bougher, S.W., et al.: Three-dimensional study of Mars upper thermosphere/ionosphere and hot oxygen corona: 2 Solar cycle, seasonal variations, and evolution over history. J. Geophys. Res. Planets 114(E11) (2009) Villanueva, G.L., Mumma, M.J., Novak, R.E., et al.: A sensitive search for organics (CH4 , CH3 OH, H2 CO, C2 H6 , C2 H2 , C2 H4 ), hydroperoxyl (HO2 ), nitrogen compounds (N2 O, NH3 , HCN) and chlorine species (HCl, CH3 Cl) on Mars using ground-based high-resolution infrared spectroscopy. Icarus 223(1), 11–27 (2013) Withers, P., Pratt, R.: An observational study of the response of the upper atmosphere of Mars to lower atmospheric dust storms. Icarus 225(1), 378–389 (2013) Zurek, R.W., Tolson, R.A., Bougher, S.W., et al.: Mars thermosphere as seen in MAVEN accelerometer data. J. Geophys. Res. Space Phys. 122(3), 3798–3814 (2017)

Chapter 8

Atmospheric Escape from Mars

Abstract Several mechanisms can be responsible for atmospheric escape. These processes can be divided into thermal escape, non-thermal escape, and impact erosion. The relative importance of each loss process depends on the planets escape velocity, its atmospheric composition and its distance from the sun. The escape occurs when molecular kinetic energy over comes gravitational energy. In other words a molecule can escape when it is moving faster than the escape velocity of its planet. Keywords Thermal · Non-thermal escape · Modeling

8.1 Earlier Measurements: PHOBOS-2 During February–March 1989, plasma was measured in-situ in the Martian tail by the Automatic Space Plasma Experiment with a Rotating Analyzer (ASPERA) and Solar Wind Plasma Experiment (TAUS) onboard Phobos 2 (Lundin et al. 1990; Rosenbauer et al. 1989). These measurements reveal abundant outflows of ionospheric ions CO2 + , CO+ , O2 + , and O+ toward the Martian tail with energy ranging from a few eV to several KeV. Lundin et al. (1989) suggested that these ions are transported in the tail either by an ion pickup process that results from the Martian boundary layer or by upward acceleration process occurring in the polar region similar to those observed above Earth’s auroral oval. Based on magnetic field measurements made by Mariner 4, Mars 2–5, Phobos 2, MGS and MAVEN, it is believed that Mars has a hybrid magnetosphere (Smith et al. 1965; Dolginov et al. 1972; Vaisberg and Smirnov 1986; Axford 1991; Möhlmann et al. 1992; Ness et al. 1999; Acuna et al. 2001; Brain et al. 2007; Connerney et al. 2015a, b)which is a combination of an intrinsic magnetic field emanating from the pole of a weak dipole and a Venus like draped interplanetary field applying in the neighbourhood of the magnetic equator beyond 2.8 Rm (Rm is equal to 3390 km). In the presence of an intrinsic magnetic field at the Martian pole the polar wind and the plasma mantle acceleration processes can provide the observed plasma population in the magneto tail under the influence of charge separation and convection electric © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_8

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fields respectively (Verigin, et al. 1991). Lammer and Bauer (1992) used a convective electric field to transport ions O2 + and CO2 + from dayside cusp region to the night side magnetosphere. Their calculations have reported a total escape rate of the order of 1022 –1023 s−1 for these ions. Haider (1995, 1996) have used charge separation electric field for the transportation of O+ from the night side ionosphere toward the plasma sheet. The total escape rate of O+ at 2300 K is obtained as ~ 3.5 × 1024 s−1 in this calculation.

8.2 Latest Measurements: MEX and MAVEN While several mechanisms can remove atmosphere from Mars over a time period, a prominent hypothesis suggests that the lack of intrinsic Earth like global magnetic dipole has allowed the solar wind to erode the early Martian atmosphere by impacting energy to the planet’s ionosphere, which subsequently flows out as ion escape. To observe the amount of atmosphere lost from Mars through the escape, ASPERA and STATIC instruments onboard MEX and MAVEN have been used respectively. ASPERA observed solar wind plasma and accelerated ions all the way down up to 270 km above the dayside planetary surface (Lundin et al. 2004). This is very deep in the ionosphere, implying direct exposure of the Martian topside atmosphere to solar wind plasma forcing. Using ASPERA data from 2007 to 2020, Nilsson et al. (2021) observed that ion escape from Mars is clearly modulated by solar activity being up to three times higher during the solar maximum than the solar minimum period. The net escape rate reaches up to about 1025 s−1 during solar maximum period. Recently STATIC instrument onboard MAVEN also observed seasonal variability of neutral escape of H and O from Mars (Rahmati et al. 2018). The escape rates as low as 3 × 1025 s−1 near the aphelion and as high as 4 × 1026 s−1 were observed near the perihelion. The O escape rates imply a much less seasonal variability with a mean value of ~ 9 × 1025 s−1 .

8.3 Thermal Escape Mechanism Thermal escape occurs when the molecular velocity due to thermal energy is sufficiently high. Thermal escape is described by Jeans escape and hydrodynamic escape mechanisms. The kinetic energy of a gas molecule depends on mass and velocity. Some molecules have much higher speed than the escape velocity. These particles will escape the atmosphere. This happens predominantly in the exosphere, where mean free path is comparable to the atmospheric scale height. The number of particles which are able to escape also depends on the molecular concentration at the exobase, which is limited by diffusion through the thermosphere. Three factors contribute to Jeans escape mechanism: (1) mass of the molecule, (2) escape velocity of the planet, and (3) heating of the upper atmosphere by solar

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radiation. The heavier molecules are likely to escape less because they have slower escape velocity than the lighter molecules. Further, a planet with larger mass have more gravity so the escape velocity tends to be greater and fewer particles will gain the energy required to escape. Therefore, the gas giant planets will retain significant amount of hydrogen, which escape more readily from Earth’s atmosphere. Finally, a distant planet from the sun has a cooler atmosphere with lower velocity and has less chance of escape. Similarly, a close planet to the sun has a hotter atmosphere with higher velocity and hence has a chance of greater escape. Under hydrodynamic condition, a large amount of thermal energy through solar radiation is absorbed by atmospheric gases. In this process the molecules are heated, they expand upwards and are further accelerated until they reach to the escape velocity. The hydrodynamic escape has been observed on exoplanets and hot Jupiter.

8.4 Non-Thermal Escape Mechanism The non-thermal escape occurs due to photochemistry or charged particle interaction. The photo dissociation can break a molecule into smaller components and provide enough kinetic energy for those components to escape. The photo ionization produces ions, which can get trapped in the planets’ magnetosphere or undergo dissociative recombination. In the first case ions may escape due to sputtering, charge exchange and polar wind processes. The sputtering escape is pronounced in absence of planetary magnetosphere, where solar wind charged particles are deflected by induced magnetic field, which mitigates the loss of atmosphere. The charge exchange escape occurs when a fast moving ion captures an electron from a slow atmospheric neutral and the new fast neutral can escape from the atmosphere, while the new slow ions will be trapped on the magnetic field lines. The atmospheric molecules can also escape from polar region on a planet with magnetosphere due to polar wind. In the polar region of a magnetosphere, the field lines are open which allow to escape the ions from the atmosphere.

8.5 Modeling of Escape Flux and Density Over the last 2–3 decades, a suit of models have been developed to simulate escape flux in the upper atmosphere of Mars (Lillis et al. 2015). First, from ground to exosphere model known as Global Circulation Models (GCMs) simulate the atmosphere (including thermosphere and ionosphere) in the fluid regime as it responds to topography, planetary rotation and solar heating and ionization (Lian et al. 2012; Millour et al. 2014, Wang et al. 2018; Bertrand et al. 2020). Second, DSMC models simulate the neutral atmosphere with macro- particles from a few scale heights below the exosphere out to several Mars radii in order to capture the physics from collisional to non-collisional regions (Valeille et al. 2009). DSMC models take part from GCMs

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near their lower boundary and are not time dependent. Last, global plasma models simulate escape flux due to interaction between the solar wind and the Martian ionosphere and exosphere using a Magneto Hydro Dynamic (MHD) approach where ions are treated kinetically and electrons are simulated as a massless charge balancing fluid (Liu et al. 2001; Ma et al. 2004, 2015; Dong et al. 2015). Global plasma models take exospheric inputs from DSMC models and thermosphere inputs from GCMs. There also exist several active modeling efforts for atmospheric/ionospheric escape mechanisms (cf. Haider (1995) , Haider (1996); Chassefiere and Leblanc 2004; Ma and Nagy 2007; Lee et al. 2015; Brecht et al. 2016, 2017; Chaffin et al. 2018; Rahmati et al. 2017, Rahmati et al. 2018).

References Acuña, M.H., et al.: Magnetic field of Mars: summary of results from the aerobraking and mapping orbits. J. Geophys. Res. Planets 106(E10), 23403–23417 (2001) Axford, W.I.: A commentary on our present understanding of the Martian magnetosphere. Planet Space Sci. 39(1–2), 167–173 (1991) Bertrand, T., Wilson, R.J., Kahre, M.A.: Simulation of the 2018 global dust storm on Mars using the NASA Ames Mars GCM: a multitracer approach. J. Geophys. Res. Planets 125(7), e2019JE006122 (2020) Brain, D.A., Lillis, R.J., Mitchell, D.L., Halekas, J.S., Lin, R.P.: Electron pitch angle distributions as indicators of magnetic field topology near Mars. J. Geophys. Res. Space Phys. 112(A9) (2007) Brecht, S.H., Ledvina, S.A., Bougher, S.W.: Ionospheric loss from Mars as predicted by hybrid particle simulations. J. Geophys. Res. Space Phys. 121(10), 10–190 (2016) Brecht, S.H., Ledvina, S.A., Jakosky, B.M.: The role of the electron temperature on ion loss from Mars. J. Geophys. Res. Space Phys. 122(8), 8375–8390 (2017) Chaffin, M.S., Chaufray, J.Y., Deighan, J., et al.: Mars H escape rates derived from MAVEN/IUVS Lyman alpha brightness measurements and their dependence on model assumptions. J. Geophys. Res. Planets 123(8), 2192–2210 (2018) Chassefière, E., Leblanc, F.: Mars atmospheric escape and evolution; interaction with the solar wind. Planet Space Sci. 52(11), 1039–1058 (2004) Connerney, J.E.P., Espley, J., Lawton, P., et al.: The MAVEN magnetic field investigation. Space Sci. Rev. 195(1), 257–291 (2015a) Connerney, J.E., Espley, J.R., DiBraccio, G.A., et al.: First results of the MAVEN magnetic field investigation. Geophys. Res. Lett. 42(21), 8819–8827 (2015b) Dolginov, S.S., Yeroshenko, Y.G., Zhuzgov, L.N.: The magnetic field in the very close neighborhood of Mars according to data from the Mars 2 and Mars 3 spacecraft. (No NASA-TM-X-66116) (1972) Dong, C., Ma, Y., Bougher, S.W., et al.: Multifluid MHD study of the solar wind interaction with Mars’ upper atmosphere during the 2015 March 8th ICME event. Geophys. Res. Lett. 42(21), 9103–9112 (2015) Haider, S.A.: O+ escape through the plasma sheet of Mars and its causative mechanism. J. Geophys. Res. Space Phys. 100(A7), 12235–12242 (1995) Haider, S.A.: High-latitude plasma transport through the Martian tail: polar wind. J. Geophys. Res. Space Phys. 101(A11), 24955–24963 (1996) Lammer, H., Bauer, S.J.: A Mars magnetic field: constraints from molecular ion escape. J. Geophys. Res. Planets 97(E12), 20925–20928 (1992) Lee, Y., Combi, M.R., Tenishev, V., et al.: A comparison of 3-D model predictions of Mars’ oxygen corona with early MAVEN IUVS observations. Geophys. Res. Lett. 42(21), 9015–9022 (2015)

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Lian, Y., Richardson, M.I., Newman, C.E., et al.: The Ashima/MIT Mars GCM and argon in the martian atmosphere. Icarus 218(2), 1043–1070 (2012) Lillis, R.J., Brain, D.A., Bougher, S.W., et al.: Characterizing atmospheric escape from Mars today and through time, with MAVEN. Space Sci. Rev. 195(1), 357–422 (2015) Liu, Y., Nagy, A.F., Gombosi, T.I., et al.: The solar wind interaction with Mars: results of threedimensional three-species MHD studies. Adv. Space Res. 27(11), 1837–1846 (2001) Lundin, R., Zakharov, A., Pellinen, R., et al.: First measurements of the ionospheric plasma escape from Mars. Nature 341(6243), 609–612 (1989) Lundin, R., Zakharov, A., Pellinen, R., et al.: ASPERA/Phobos measurements of the ion outflow from the Martian ionosphere. Geophys. Res. Lett. 17(6), 873–876 (1990) Lundin, R., Barabash, S., Andersson, H., et al.: Solar wind-induced atmospheric erosion at Mars: first results from ASPERA-3 on Mars Express. Science 305(5692), 1933–1936 (2004) Ma, Y.J., Russell, C.T., Fang, X., et al.: MHD model results of solar wind interaction with Mars and comparison with MAVEN plasma observations. Geophys. Res. Lett. 42(21), 9113–9120 (2015) Ma, Y.J., Nagy, A.F.: Ion escape fluxes from Mars. Geophys. Res. Lett. 34(8) (2007) Ma, Y., Nagy, A.F., Sokolov, I.V., Hansen, K.C.: Three-dimensional, multispecies, high spatial resolution MHD studies of the solar wind interaction with Mars. J. Geophys. Res. Space Phys. 109(A7) (2004) Millour, E., Forget, F., Spiga, A., et al.: A new Mars climate database v5 1. In: The Fifth International Workshop on the Mars Atmosphere: Modelling and Observations, pp. id-1301 (2014) Möhlmann, D., Sauer, K., Roatsch, T., et al.: Magnetic field environment of Mars as studied by Phobos-2. Adv. Space Res. 12(9), 221–229 (1992) Ness, N.F., Acuña, M.H., Connerney, J., et al.: MGS magnetic fields and electron reflectometer investigation: discovery of paleomagnetic fields due to crustal remanence. Adv. Space Res. 23(11), 1879–1886 (1999) Nilsson, H., Zhang, Q., Wieser, G.S., et al.: Solar cycle variation of ion escape from Mars. Icarus 114610 Rahmati, A., Larson, D.E., Cravens, T.E.: MAVEN measured oxygen and hydrogen pickup ions: probing the Martian exosphere and neutral escape. J. Geophys. Res. Space Phys. 122(3), 3689– 3706 (2017) Rahmati, A., Larson, D.E., Cravens, T.E.: Seasonal variability of neutral escape from Mars as derived from MAVEN pickup ion observations. J. Geophys. Res. Planets 123(5), 1192–1202 (2018) Rosenbauer, H., Shutte, N., Apathy, I., et al.: Ions of Martian origin and plasma sheet in the Martian magnetosphere: initial results of the TAUS experiment. Nature 341(6243), 612–614 (1989) Smith, E.J., Davis, L., Jr., Coleman, P.J., Jr., Jones, D.E.: Magnetic field measurements near Mars. Science 149(3689), 1241–1242 (1965) Vaisberg, O., Smirnov, V.: The martian magnetotail. Adv. Space. Res. 6(1), 301–314 (1986) Valeille, A., Combi, M.R., Bougher, S.W. et al (2009) Three-dimensional study of Mars upper thermosphere/ionosphere and hot oxygen corona: 2 Solar cycle, seasonal variations, and evolution over history. J. Geophys. Res. Planets 114(E11) (2009) Verigin, M.I., Shutte, N.M., Galeev, A.A., et al.: Ions of planetary origin in the Martian magnetosphere (Phobos 2/TAUS experiment). Planet Space Sci. 39(1–2), 131–137 (1991) Wang, C., Forget, F., Bertrand, T., et al.: Parameterization of rocket dust storms on Mars in the LMD Martian GCM: modeling details and validation. J. Geophys. Res. Planets 123(4), 982–1000 (2018)

Chapter 9

Upper Ionosphere of Mars

Abstract The upper ionosphere of Mars (above 100 km) is mainly produced by X-rays (10–100 Å) and solar EUV radiation (100–1026 Å) incident on the top of the atmosphere. The E and F regions are produced at about 110 and 130 km due to X-rays and solar EUV radiation respectively. Above 200 km, the exosphere is formed where collision is almost negligible and mean free path becomes very large. The ions can escape from the upper ionosphere when the kinetic energy of the ion is faster than the gravitational binding energy. The solar wind also impact with the upper ionosphere because Mars has an induced magnetosphere. The upper ionosphere is highly disturbed, when a large solar flare, Coronal Mass Ejections (CMEs) and Solar Energetic Particles (SEPs) impacted with the Mars’ atmosphere. The upper ionosphere of Mars is also changing with minimum to maximum solar activity period. Keywords Ionosphere · Electron density · Measurements

9.1 Photochemical Equilibrium Region The differential form of the continuity equation is described as follows: ∂n s = ps − n s l s ∂t

(9.1)

where ns is the density of ion s, ps and ls are the production rates and loss coefficients of ion s. Under steady state condition ∂n s = 0, then ∂t ps = n s l s

(9.2)

Equation (9.2) is known as photochemical equation, which can be used in the photochemical region of the upper ionosphere of Mars. This equation is also © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_9

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described in Chaps. 2 and 7. This region occurs in the upper ionosphere of Mars between 100 and 200 km.

9.2 Diffusive Equilibrium Region In Sect. 2.2 the vertical momentum equation is defined as: ∂n s + ∇ · (n s · u→) = 0 ∂t

(9.3)

where t is time and u→ is diffusion velocity in the vertical direction. Under steady state s = 0, then condition, ∂n ∂t ∇ · (n s · u→) = 0 ns · u→ = constant This is known as diffusive equilibrium condition, where flux is constant. In the top of the ionosphere (above 200 km) ambipolar diffusion is important. In this region chemical life time (τ = l−1 ) is very large than the diffusion time (τ m = H 2 /D), where l is the loss coefficient, H and D is scale height and diffusion coefficient respectively (St. Maurice and Schunk (1977).

9.3 Plasma Transport Processes The upper ionosphere of Mars is quite variable, which seems to violate diffusive equilibrium. In presence of purely induced horizontal magnetic field the plasma loss is due to horizontal divergence of ion velocities. In presence of vertical magnetic field an upward flow of plasma seems to play an important role in the upper ionosphere. The Martian ionosphere is also affected due to drift velocity, which can be represented in terms of zonal, meridional and vertical winds of the neutral atmosphere. It varies from 10 m/s to 50 m/s in the upper ionosphere of Mars (Majeed et al. 2012). The drift velocity reduces the electron density insignificantly in comparison to that reduced by horizontal plasma transport velocity (Haider et al. 2023). In the plasma transport processes ions and electrons are moving by electric fields. The resulting motions and electric currents depend on the magnetic field and the collision frequencies, which determine the mobility and electrical conductivity of the charged particles. The charged particles can be moved by neutral air wind. The daily temperature changes in the thermosphere affect the charged particles and the neutral air. Hence the plasma takes part in the thermal expansion and contraction of the atmosphere. The plasma tends to diffuse under the action of gravity and of

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gradients in its own partial pressure. The electrical forces among ions and electrons tend to keep them together so that both kinds of particles diffuse at the same speed.

9.4 Early Measurements: Mariner 4, 6, 7 and 9, Mars 4, 5 and Viking 1, 2 The first measurement of the Martian ionosphere was carried out by the radio occultation experiment aboard Mariner 4 on 15 July 1965 (Kliore et al. 1965). A dayside ionosphere from 90 to 250 km with a peak value of about 105 cm−3 at an altitude of 125 km and a topside scale height of 29 km was observed. The nighttime ionosphere could not be detected as the sensitivity of this instrument was limited to an electron density of 5 × 103 cm−3 (cf. Verigin et al. 1991). Mariner 4 was followed by Mariner 6, 7, and 9 and Viking 1 and 2 (Rasool and Stewart 1971). Mariner 9 provided a long period of observations and generated two sets of data, one during November–December 1972 and the other during May– June 1972. The nighttime ionosphere of Mars has been observed first from radio occultation experiment onboard Mars 4 and 5 on 10 and 18 February, 1974 (Savich and Samovol 1976). The electron density profiles observed by Mars 4 and 5 are shown in Fig. 24.1. These measurements showed a clear peak at altitude ~ 150 km due to precipitation of solar wind electron (Haider et al. 2007, 2009) The dayside and nightside ionosphere of Mars were also measured by Viking 1 and 2. The peaks of the electron density profiles for the nightside and dayside ionosphere are found at about 150 km and 130 km respectively. The major ion densities O2 + , CO2 + and O+ are measured by Viking 1 and 2 in the daytime atmosphere. The ion densities are not measured in the nighttime ionosphere of Mars (Hanson et al. 1977). Several profiles in the nighttime ionosphere of Mars are very noisy and do not show any significant peak (Zhang et al. 1990). This may be due to the fact that the electron density in the nighttime ionosphere of Mars is very low, which were not possible to detect by RO experiment.

9.5 Latest Measurements: MGS, MEX, MARSIS and MAVEN The RO experiment on MGS obtained 5600 electron density profiles during 24 December 1998 to 9 June 2005. Each profiles observed a primary peak at about ~ 135 km. There is also a secondary peak in the Martian ionosphere at about 105– 120 km (Bougher et al. 2001). These primary and secondary peaks are known as F and E layers respectively. The MEX in addition to RO experiment has a low frequency radar known as MARSIS experiment, which transmits radio waves from 0.1 to 5.4 MHz and thus can provide topside ionograms. These ionograms can give

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Fig. 9.1 Electron density profiles observed by ROSE onboard MAVEN (from Withers et al. 2020)

electron density profiles between F peak and altitudes as high as 500 km in the top side ionosphere. Apart from E and F layers, MEX also observed a meteoric layer for the first time at altitudes between 80 and 100 km (Pätzold et al. 2005; Withers et al. 2008; Haider et al. 2013; Pandya and Haider 2014). Recently ROSE onboard MAVEN has observed an electron density profile on 10 October 2017. This profile has two layers at altitudes ~ 130 and ~ 80 km due to photoionization and meteoroid impact ionizations respectively (Haider et al. 2023). In Fig. 9.1 the E layer is not visible clearly at about ~ 110 km due to X-ray impact ionization (Thirupathaiah et al. 2019).

9.6 Solar Zenith Angle Dependence of Mars’ Ionosphere The dayside ionosphere of Mars produces E and F layers due to impact of X-ray and EUV radiations at altitudes ~ 110 km and ~ 130 km respectively. Thirupathaiah et al. (2019) have analyzed 280 electron density profiles with an apparent solar flare response in 32 profiles. Among 32 flare profiles 10, 12 and 10 profiles were strongly perturbed by X, M and C class flares respectively. In Fig. 9.2a, b the variations of the E-peak electron densities and E- peak altitudes of 32 flare profiles are shown with SZA. The peak electron densities are decreasing with increasing SZA. The peak heights of the electron densities are increasing with increasing SZA. The zenith angle dependence of the peak electron densities can be fitted by a Chapman formula N m = N o cos0.5 θ distribution for θ < 90°, where θ is a SZA, N o is a peak electron density at SZA = 0°, and N m is a peak electron density at SZA > 0°. The altitudes of the peak electron densities can be fitted by hm = ho + H ln sec θ, where hm is a peak height, H is a neutral scale height ~ 10 km (Zhang et al. 1990) and ho is ~ 95 km for E region ionosphere.

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(a)

(b)

Fig. 9.2 Variation of peak electron densities (a) and peak altitudes (b) of 32 flare profiles as a function of SZA. The different colour symbols show the type of the flares. The solid lines represent the Chapman fit to the peak electron densities and peak altitudes. (from Thirupathaiah et al. 2019)

References Bougher, S.W., Engel, S., Hinson, D.P., Forbes, J.M.: Mars global surveyor radio science electron density profiles: neutral atmosphere implications. Geophys. Res. Lett. 28(16), 3091–3094 (2001) Haider, S.A., Prasad, D., Shah, S.: The magnetically controlled ionopause boundary observed by LPW onboard MAVEN within magnetic pile-up region of Mars. In Press, Icarus (2022) Haider, S.A., Singh, V., Choksi, V.R., Maguire, W.C., Verigin, M.I.: Calculated densities of H3 O+ (H2 O)n, NO2 − (H2 O)n, CO3 − (H2 O) n and electron in the nighttime ionosphere of Mars: Impact of solar wind electron and galactic cosmic rays. J. Geophys. Res. Space Phys. 112(A12) (2007) Haider, S.A., Abdu, M.A., Batista, I.S., et al.: D, E, and F layers in the daytime at high-latitude terminator ionosphere of Mars: comparison with earth’s ionosphere using COSMIC data. J. Geophys. Res. Space Phys. 114(A3) (2009)

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Haider, S.A., Pandya, B.M., Molina-Cuberos, G.J.: Nighttime ionosphere caused by meteoroid ablation and solar wind electron-proton-hydrogen impact on Mars: MEX observation and modeling. J. Geophys. Res. Space Phys. 118(10), 6786–6794 (2013) Haider, S.A., Prasad, K.D., Shah, S.Y.: The magnetically controlled ionopause boundary observed by LPW onboard MAVEN within magnetic pile-up region of Mars. Icarus 394, 115423 (2023) Hanson, W.B., Sanatani, S., Zuccaro, D.R.: The martian ionosphere as observed by the Viking retarding potential analyzers. J. Geophys. Res. 82(28), 4351–4363 (1977) Kliore, A., Cain, D.L., Levy, G.S.D., et al.: Occultation experiment: results of the first direct measurement of Mars’s atmosphere and ionosphere. Science 149(3689), 1243–1248 (1965) Maurice, J.P., Schunk, R.W.: Diffusion and heat flow equations for the mid-latitude topside ionosphere. Planet Space Sci 25(10), 907–920 (1977) Majeed, T., Bougher, S.W., Haider, S.A.: Plasma transport processes in the topside Martian ionosphere. Advances in Geosciences: Volume 30: Planetary Science (PS) and Solar and Terrestrial Science (ST), 57–68 (2012) Pätzold, M., Tellmann, S., Hausler, B., et al.: A sporadic third layer in the ionosphere of Mars. Science 310(5749), 837–839 (2005) Rasool, S.I., Stewart, R.W.: Results and interpretation of the S-band occultation experiments on Mars and Venus. J Atmos Sci 28(6), 869–878 (1971) Savich, N.A., Samovol, V.A.: The night-time ionosphere of Mars from Mars 4 and Mars 5 dualfrequency radio occultation measurements. Space Res. XVI, 1009–1011 (1976) Thirupathaiah, P., Shah, S.Y., Haider, S.A.: Characteristics of solar X-ray flares and their effects on the ionosphere and human exploration to Mars: MGS radio science observations. Icarus 330, 60–74 (2019) Verigin, M.I., Shutte, N.M., Galeev, A.A., et al.: Ions of planetary origin in the Martian magnetosphere (Phobos 2/TAUS experiment). Planet Space Sci 39(1–2), 131–137 (1991) Withers, P., Mendillo, M., Hinson, D.P., Cahoy, K.: Physical characteristics and occurrence rates of meteoric plasma layers detected in the Martian ionosphere by the Mars global surveyor radio science experiment. J. Geophys. Res. Space Phys. 113(A12) (2008) Withers, P., et al.: The MAVEN radio occultation science experiment (ROSE). Space Sci. Rev. 216(4), 61 (2020) Zhang, M.H.G., Luhmann, J.G., Kliore, A.J.: An observational study of the nightside ionospheres of Mars and Venus with radio occultation methods. J. Geophys. Res. Space Phys. 95(A10), 17095–17102 (1990)

Chapter 10

Models of the Martian Ionosphere

Abstract In this Chapter we have described important models which have been used to study the Martian ionosphere and give several examples of their usage (Haider and Mahajan in Space Sci Rev 182:19–84, 2014). We start from the Boltzmann equation and continue to a macroscopic description of the plasma using continuity, momentum and energy equations in Sect. 10.1. MHD and Hybrid models are selfconsistent Mars-solar wind plasma approaches which are given in Sects. 10.2 and 10.3 respectively. The two-stream method is described in Sect. 10.4. In Sect. 10.5 Monte Carlo modeling illustrates a kinetic model. Finally Analytical Yield Spectrum (AYS) method is described in Sect. 10.6. This approach is based on Monte Carlo method. In Sects. 10.7 and 10.8 we have described energy loss and meteoroid ablation model respectively. Keywords Modeling instruments · Ionosphere · Atmosphere

10.1 Boltzmann Transport Model In the Boltzmann equation the properties of the plasma are described by a 6-D distribution function, f (r, ν). The distribution function gives the plasma number density at a given time in a space volume, d 3 r, and velocity space volume, d 3 v. The Boltzmann equation can be written in the following form (Haider et al. 2011) ∂ F→s + v→ · ∇r F→s + a · ∇v F→s − Q(→ r , v→, t) − ∂t

(

δ F→s δt

) =0

(10.1)

Coll.

where ∂ F→s /∂t represents local variation of distribution function F→s of species s with time t, ∇r is the space derivative, ∇v is the velocity derivative, ν→ is the velocity, and a is the acceleration. Therefore, the second term on the left-hand side describes the change in F→ s due to spatial gradients, and the third term describes the presence of external forces, Q is the source term for production rates obtained from solar radiation, and (∂ F→s /∂t)Coll represents chemical production /loss rates. Under a number © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_10

69

70

10 Models of the Martian Ionosphere

of simplifying assumptions the transport Eq. (10.1) was solved numerically for the photoelectron distribution in the upper atmosphere of Mars (Mantas and Hanson 1979, 1985, 1987). The Boltzmann equation is rarely used in the form of Eq. (10.1). For space and ionospheric physics it is represented as fluid or conservation equations. These equations can be derived by taking velocity moments of the Boltzmann equation. The zero, first, second, and third velocity moments of the distribution function, F sp (r,v,t), became the density (ns ), flow velocity (us ), pressure tensor, and heat flux tensor, respectively. The zero moment of the Boltzmann equation yields the continuity equation, the first moment yields the momentum equation, the second moment yields the energy equation, and the third moment yields the heat flow equation as follows:

10.1.1 Continuity and Momentum Equations ∂n s + ∇ · (n s · u→s ) = Ps − L s ∂t

(10.2)

where Ps is the production rates of species s due to photon impact, photoelectron impact and including primary production by photoionization and photo electron impact ionization chemical reactions. The L S is the loss rate of species s due to chemistry of different chemical reactions. This equation is solved by several investigators to study the Martian ionosphere (Fox 1993), Fox (2009); Fox and Yeager 2006, Fox and Yeager 2009; Haider et al. 2010 and references therein). The momentum equation for species s in presence of magnetic and electric field can be written as [ ns m s

] ( ) ∂ u→s + u→s · ∇ u→s = −∇ ps + n s es · E→ + u→s × B→ + n s m s · g→ ∂t ∑ ( ) ( ) νs j u→s − u→ j + ps m s u→s − u→ j − ns m s

(10.3)

j

where g is the acceleration due to gravity and ν sj is the momentum transfer collision frequency between species s and j. The species s denotes electrons or ions and ms is the mass of charged particles, E and B are the electric and magnetic fields, respectively, ps is the pressure (= ns K B T s , where K B and T s are Boltzmann constant and temperature, respectively), and un is the speed of the neutral. The momentum equation for electrons can be rearranged to find the electric field as given below: ∇ pe J→ × B→ − + η J→ − m e νen (→ E→ = −→ u × B→ + u − u→n ) en e en e

(10.4)

where J is the current density, pe is the pressure due to electrons, η is the resistivity, me is the mass of electrons, ne is electron density, ν en is electron-neutral collision

10.1 Boltzmann Transport Model

71

frequency, and u is average plasma flow speed. Equations (10.3) and (10.4) can be combined to give a single-fluid momentum equation as follows: ) ∂ u→ + u→ · ∇ u→ = J→ × B→ − ∇( pe − pi ) − ρ g→ − ρνin (→ ρ u − un ) ∂t (

where ρ is the mass density, pi and ν in are ion pressure and ion-neutral frequency, respectively. By using Ampere’s law the J × B force term can be split into a magnetic pressure gradient force term and a magnetic tension force term. By neglecting the lefthand term and solving for plasma flow speed u, one can get the ambipolar diffusion equation (Chen et al. 1978; Shinagawa and Cravens 1989).

10.1.2 Energy and Heat Flow Equations The energy equation is obtained from the second moment equation of the Boltzmann equation as follows: ( ) 3 ∂ Ts ∂ Ts ∂ Ks = Qs − L s k B ns − 2 ∂t ∂z ∂z

(10.5)

where Qs and Ls are heating and cooling rates, respectively, and Ks is the coefficient of thermal conductivity. This energy equation leaves out dynamical terms such as heat advection, which is less important for electrons than heat conduction and local heating and cooling. In general, one needs yet another moment equation for heat flow, but it is approximated by the heat conduction expression in the left-hand side by the second term. Chen et al. (1978) and Rohrbaugh et al. (1979) have solved energy equations to study ion and electron temperatures in the dayside ionosphere of Mars. In the magnetized plasma the momentum equation includes the magnetic field. The magnetic field can be derived from the so-called magnetic induction equation as given below: ( ) ∂ B→ = ∇ × u→s × B→ − ∇ × ∂t

(

η ∇ × B→ μ0

) (10.6)

The first term on the right-hand side of Eq. (10.6) is the magnetic convection term and the second term is the magnetic diffusion term.

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10 Models of the Martian Ionosphere

10.2 Magneto-Hydrodynamic (MHD) Model In MHD model plasma is considered as a fluid. The ion velocity distribution is assumed to be Maxwellian. One can adopt various approximations, which lead to (1) ideal MHD, (2) multispecies MHD, (3) Hall MHD, and (4) multi-fluid MHD models. All these MHD models have been used to study solar wind interaction with non-magnetized solar system objects such as Mars, Venus, and Saturn’s moon Titan (cf. Nagy et al. 2004; Ma et al. 2004; Kallio et al. 2010 and references therein). The ideal MHD is a single fluid model where the mass density (ρ), the bulk velocity (u), the scalar thermal pressure (p), and the internal energy density (ε) obey the continuity, the momentum, and the energy equations as well as the Faraday’s law as given below: Continuity equation ∂ρ + ∇ · (ρ u→) = 0 ∂t

(10.7)

∂ρ u→ + ∇ · (ρ u→u→) = J→ × B→ − ∇. p ∂t

(10.8)

∂ε + ∇ · (εu→) = −ρ∇ · u→ ∂t

(10.9)

∂ B→ = −∇ × E→ ∂t

(10.10)

Momentum equation

Energy equation

Faraday’s law

The thermal pressure is related to the internal energy density as p = (γ − 1) ε, where γ is the specific heat ratio. The different MHD models use different forms of Ohm’s law. The ideal MHD uses the ideal Ohm’s law (E = − u × B) (the resistive MHD model includes resistivity, η, also in Ohm’slaw). The Hall MHD model includes the Hall term (~ J × B) in Ohm’s law. The multispecies MHD model includes several ion species. In this model all ion species have their own continuity equation, but the model includes only one momentum equation. Finally, the multifluid MHD models include the continuity, momentum, and energy equations for each ion species. Various MHD models are developed into one, two, and three dimensions to study the dynamics of the Martian upper ionosphere (Shinagawa and Cravens 1989; Shinagawa and Bougher 1999; Ma et al. 2004). MHD models provide the capability to perform a high resolution threedimensional simulation of the solar wind–planet interaction, which contains both

10.3 Hybrid Model

73

solar wind and self-consistent ionosphere. This approach has some limitations because it uses a simplified approximation. In single-fluid and multi-species MHD models all ion species are assumed to have the same bulk velocity. In reality, this is not the case everywhere on Mars. The pickup ions O+ have different speeds and directions from the solar wind protons (cf. Kallio et al. 2010). Furthermore, all MHD models assume a Maxwellian velocity distribution function. The velocity distribution of pickup ions O+ can be highly non-Maxwellian on Mars. The physical process of a non-Maxwellian ion is not included in the MHD model.

10.3 Hybrid Model In the absence of magnetosphere solar wind proton interact with the planetary neutral constituents to ionize and pick-up ion particles along with solar wind protons. The physics of the interaction of solar wind protons and atmospheric constituents has been studied by Hybrid model below 500 km where suitable boundary conditions are applied to solve the atmospheric ionizatio problems (Shinagawa and Cravens 1989; Shinagawa 1996). In the Hybrid model, the accelerated ions and electron obey Lorentz force (Kallio and Janhunen 2002) as given below: mi

d dt

(

d Xi dt

)

→ = e · [ E→ + U→i × B]

(10.11)

where i represents ion with E→ and B→ are electric and magnetic fields respectively, e is electron charge, Xi is the instantaneous position of the ions with bulk velocity Ui and for electron the bulk velocity is Ue and ionization frequency is νi then, above equation is reduced to d mi dt

(

d Xi dt

)

= e · [→ vi − U→e ] × B→

(10.12)

The electric field E and magnetic field B are calculated from Eqs. (10.4) and (10.10) respectively. In MHD and hybrid models, electrons are modeled in a relatively simplified manner. Typically, electrons are assumed to be a single fluid with a Maxwellian velocity distribution function. In addition, the electrons are assumed to be adiabatic or isothermal flows. In practice, several electron species exist in the Martian ionosphere, for example, thermal electrons, photoelectrons, and electrons originating from the magneto sheath. Collision of electrons results in ion sources and losses and changes in the electron velocity distribution when high-energy electrons cascade to lower energies. The two-stream method for electron collision is given below.

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10 Models of the Martian Ionosphere

10.4 Two Stream Method Hanson and Johnson (1961) and Hanson (1963) discussed the photoelectron, ion and neutral temperature in the Martian ionosphere above 300 km due to impact of solar EUV between energy ranging from few eV to 100 eV (Dalgarno et al. 1963; Bank and Nagy 1970; Willmore 1964). The photoelectrons below 300 km in the Martian atmosphere lose their energy due to elastic and inelastic collision with neutrals whose scattering cross sections are σe (ε) and σa (ε) respectively between energy range ε and ε + dε (Banks et al. 1974; Stamnes and Swanson 1981). The elastic collision of photoelectrons with atmospheric neutrals can convert downward moving photoelectrons into upward and vice-versa. The fraction of these electrons backscattered by ith atmospheric species is pei (ε). The two-stream model is deriving transport processes in both elastic and inelastic collision with photoelectron flux φ (ε, s) of s-coordinate, which can be separated into two components: (1) the component along geomagnetic field lines away from the planet surface represented by φ + (ε, s) i.e., upwards and (2) the component along geomagnetic field lines towards the planet surface represented by φ − (ε, s) i.e., downwards. The two stream method is described as given below: ∑ ] [ dφ + (ε, s) =− n i (s) σai + pei σei ·φ + (ε, s) ds i ∑ Q(ε, s) − + Q + (ε, s) n i (s) pei σei · φ (ε, s) + + 2 i ∑ ] [ dφ − (ε, s) =− n i (s) σai + pei σei ·φ − (ε, s) ds i ∑ Q(ε, s) − + Q − (ε, s) n i (s) pei σei · φ (ε, s) + + 2 i

(10.13)

(10.14)

where Q± is the photoelectron production due to cascading from higher-energy photoelectrons undergoing inelastic collisions; Q is the photoelectron production rate due to direct ionization process; s is the distance along the magnetic field lines; and nk is neutral density. This method has been developed by Banks and Nagy (1970) to calculate photoelectron upward/downward flux and production rates in Earth’s ionosphere. Gan et al. (1990) used this method in the calculation of the upper ionosphere of Venus to study the ionopause boundary layer. Chen et al. (1978) and Haider et al. (1992) modified this model for the calculation of ion production rates on Mars. Later, production rates were used in the continuity and momentum equations for the study of the Martian atmosphere.

10.6 AYS Approach

75

10.5 Monte Carlo Model In the Monte Carlo model, ions or electrons are modelled as particles. This method is useful for obtaining numerical solutions to problems that are too complicated to solve analytically. In this method a real number between 0 and 1 is generated to determine whether a collision takes place or not. If not the amount of energy lost through Coulomb losses to the ambient electrons is calculated from Butler-Buckingham formula (Dalgarno et al. 1963) and is added to accumulated energy loss. If the collision with atmospheric gases occurred a further decision is made whether the collision is elastic or inelastic. For the elastic collision, the scattering angle calculation was carried out by Poter (1978). If the scattering event is inelastic then the states which are excited and ionized were calculated from Jackman et al. (1977). Ip (1988) calculated the number densities of hot oxygen atoms at Mars and Venus using one-dimensional Monte Carlo method. Kallio and Barabash (2001) have developed a three-dimensional Monte Carlo model to study the effects of fast hydrogen atoms into the Martian ionosphere. These atoms are produced due to charge exchange between solar wind proton sand hydrogen coronas. The hydrogen corona has the same energy and move in the same direction as protons move before the collisions. In this process solar wind protons and hydrogen atoms are combined (H+ -H) and penetrate deeper into the Martian atmosphere. The H+ -H lose their energy at low altitude after collisions with atmospheric gases. Haider et al. (2002) used this model to study the Martian ionosphere caused by electron-proton-hydrogen atoms. Using Monte Carlo method AYS was developed by Green and his co-workers (Green et al. 1977; Jackman and Green 1979; Singhal et al. 1980). AYS is defined below.

10.6 AYS Approach The AYS model was initially generated by Singhal et al. (1980) and Singhal and Green (1981) using a Monte Carlo method. These authors have fitted yield spectra analytically. Later AYS model has been used and extended by several investigators (Haider and Singhal 1983; Singhal and Haider 1984; Haider and Bhardwaj 2005; Haider et al. 2011; Haider and Mahajan 2014; Pandya and Haider 2014; Thirupathaiah et al. 2019; Shah et al. 2021; Haider and Masoom 2019). In this method mono energetic electrons of energy Eo are introduced in a gas medium. The energy of secondary or tertiary electrons and its position are calculated at that time when the primary electrons excite/ionize the atmospheric constituents. In this process an yield spectrum function was produced. This function calculates the yield of any state in the mixture of gases. Later this function was fitted analytically. The yield spectrum is presented in terms of two, three, four and five dimensionals (Green et al. 1977; Singhal et al. 1980; Haider and Singhal 1983; Singhal and Haider 1984) as given below.

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10 Models of the Martian Ionosphere

10.6.1 Two-Dimensional Yield Spectra The two-dimensional yield spectrum U(E,Eo ) in unit (eV)−1 is defined as U (E, E 0 ) = σT (E) f (E, E 0 ) = N (E)/ΔE

(10.15)

where σT (E) is the total inelastic cross section and N (E) is the number of electrons in the bin centered at E after one bin has emptied and before the next lower non empty bin of width Δ E centered at E is considered, f (E, Eo ) is the equilibrium flux or degradation spectrum of Spencer and Fano (1954). The analytic form of Eq. (10.15) is given below (Singhal and Haider 1984, 1986; Haider 1986): U (E, E o ) = Ua (E, E o ) H (E o − E − E min ) + δ(E o − E)

(10.16)

where H is Heaviside function with Em the minimum threshold of the states considered, δ (Eo −E) is the Dirac delta function, E is the energy of secondary electrons, and Eo is the energy of incident electrons in eV. Ua (E, Eo ) can be expressed as Ua (E, E o ) = R0 + R1 Y + R2 Y 2 Y =

(E o /1000)0.585 E +1

(10.17)

(10.18)

where R0 , R1 and R2 are adjustable parameters for different gases of Mars (e.g., Green et al. 1977). This model is useful where the field lines are horizontal and photoelectrons/electrons lose their energy where they are produced.

10.6.2 Three-Dimensional Yield Spectra The energy of incident electrons Eo is degraded completely and results to the total number of electrons N(E,Z) exist in the spatial interval ΔZ with Z as centre and in the energy range ΔE with E as centre after interaction. The two dimensional yield spectra (10.17) is modified into three dimensional yield spectra U (E, Z, Eo ) in unit eV−1 (gm/cm2 )−1 . Three dimensional yield spectra are used in presence of vertical magnetic field where vertical transport of the electrons is allowed as: U (E, Z , E 0 ) =

N (E, Z ) ΔE ΔZ

(10.19)

The analytic form of three dimensional yield spectra is given below (Haider and Singhal 1983; Singhal and Haider 1984).

10.6 AYS Approach

77

U (E, Z , E o ) = Ua (E, Z , E o ) H (E min − E − E o ) − δ (E o − E) D(Z , E o ) (10.20) where δ(E o − E) and H (E min − E − E o ) are Dirac Delta and Heaviside functions, respectively and D (Z , E o ) is a source term as given below: D(Z , E o ) =

d11 χ d12 / exp (− R d2 )

(10.21)

In Eq. (10.20) E, Eo and Emin are same as given in Eq. (10.16). The d2 , d11 , d12 , and χ are adjustable parameters. R is a scale factor. These are analytical parameters which are given by Green et al. (1977) and Haider and Singhal (1983). Ua (E, Z, Eo ) is an analytical yield spectra which can be expressed as given below: Ua (E, Z , E o ) = A (Z , E o ) + B1 (Z , E o ) (E R ) B2 + C1 (Z , E o ) (E R )C2 (10.22) The parameters A, B1 , B2 , C1 , C2 and ER are given by Haider and Singhal (1983). Three dimensional yield spectra is used for the calculation of ion production rates in presence of crustal magnetic field.

10.6.3 Four-Dimensional Yield Spectra The four dimensional yield spectra depends on the total number of inelastic collisions N (E, r, Z), which occur in the spatial interval ΔZ and radial interval ΔS with centre (r, z). The symbol r is a radial distance which can be calculated from the formula r 2 = x 2 + y 2 , where x and y are the cylindrical coordinates. The unit of four dimensional yield spectra is eV−1 (gm/cm2)−3 (Jackman and Green 1979). The four dimensional yield spectra is defined as U (E, r, z, Eo) = [( ΔS =

Δr r+ 2

N (E, r, Z ) ΔE · ΔZ · ΔS

)2

) ] ( Δr 2 − r− 2

(10.23)

(10.24)

The four dimensional yield spectra is integrated over radial distance r to obtain three dimensional yield spectra. But this integration is only applicable if spatial distribution of plasma is not required in electron motion. The gyro radius of the plasma should not be very large. The longitudinal distribution of plasma cannot be studied from three dimensional yield spectra. In auroral studies, longitudinal distance of electron impact along with the radial distance is an important parameter to be added in the three dimensional yield spectra.

78

10 Models of the Martian Ionosphere

10.6.4 Five-Dimensional Yield Spectra The impact of energetic photo electrons in the planetary atmosphere suffer elastic scattering due to Coulomb interaction and produce spectra comprising of a continuous radiation known as bremsstrahlung spectra (Luhmann 1977). In auroal events we need to study distribution of Polar angle (θ) associated with electron spectra. This requires another degree of freedom in four dimensional yield spectra to result into five dimensional yield spectra whose analytic form has been represented by Singhal (1980) and Riewe and Green (1978) as U (E, r, z, θ, E o ) =

2 ∑ Ai · G i (r, z) · Q i 3 R i=1

(10.25)

where Qi is a normalized angular distribution function represented by ∫π Q i dθ = 1

(10.26)

0

10.7 Energy Loss Model In the energy loss method, a charge particle losses its kinetic energy or it is deflected from its original path involving four principal type of interactions: (1) Inelastic collision with bound atomic electrons are usually the predominant mechanism by which a charged particle losses kinetic energy in an absorber. As a result of such collision, one or more atomic electrons experience a transition to an excited state or to an ionization state, (2) Inelastic collision with a nucleus experiences a deflection. In some but not all such deflections, a quantum of radiation is emitted (bremsstrahlung) and a corresponding amount of kinetic energy is lost by the colliding particles, (3) In inelastic scattering the incident particle is deflected but does not radiate, nor does it excite the nucleus. The incident particle losses kinetic energy required for conservation of momentum between the two particles. (4) An incident charge particle may be elastically deflected in the field of the atomic electron of struck atom. Energy and momentum are conserved, and energy transfer is generally less than the lowest excitation potential of the electrons, so that the interaction is really with the atom as a whole. Such collisions are significant for the case of very low energy (< 100 eV) incident electrons. In the absorbing material, a moving particle is slowed down and finally brought to at rest by the combined action of all four of these elastic and inelastic collisions processes. From collision theory, one can obtain the probabilities of any particular change of the direction of motion of incident particle. After the first collision, these probabilities can be applied a second collision, then to a third etc.

10.8 Meteoroid Ablation Model

79

The basic formula for energy loss per cm path into the material media by a particle traveling with the velocity ‘v’ and undergoing inelastic collisions is written as given by Evans (1995); Haider et al. (2007). (

dE dh

)

∫ ( ∫ ) e4 m o v2 E 2 ( ) −β = 2π N Z ln m o v2 I 2 1 − β2

(10.27)

In Eq. (10.27), by substituting mo v2 = β 2 mo c2 and E + mo c2 = mo c2 (1−β 2 ) , we get

−0.5

(

dE dh

) = 4π

moc ro2 2 β

2

∫

(

E + m o c2 I

N Z ln β

)(

E m o c2

) 21

1 − β2 2

∫ (10.28)

where I is mean ionization potential, N is the neutral density, ro is the classical electron radius with 4π ro 2 = 1.0 × 10–24 cm2 /electron, β 2 = (v/c)2 = 1 − [(E/mo c2 ) + 1] −2 , mo c2 = 0.51 meV, Z is the mean atomic number, and c is the velocity of light. Using Eq. (10.28) the ion production rate due to impact of GCR at height h and solar Zenith angle χ is given below: J(h, χ ) =

2π Q

∫

∞

(

E

) dE F(χ , E)d E dh

(10.29)

where Q = 35 eV is the energy required for the formation of an electron ion pair, χ is Solar Zenith Angle, F is the total differential flux of GCR being expressed in cm−2 s−1 GeV−1 ster−1 at height h.

10.8 Meteoroid Ablation Model Meteoroids entering into Martian atmosphere interact with the gases and produce fusion, evaporation and sputtering on the surface of meteoroids which in turn deposits metallic ions and electrons in the ionosphere. The ablation of these meteoroids in the planetary atmosphere form ionospheric layers (Hughes 1978; Pesnell and Grabosky 2000; Mollina-Cuberos et al. 2003). Whipple (1950, 1951) first described interaction and heat balance between micro meteoroids and atmospheric constituents. Later meteoroid ablation and ionization theory was studied by several investigators (Öpik 1958; Adolfsson et al. 1996; Lebedinets et al. 1973; McKinley 1961; Ceplecha et al. 1998; Fisher et al. 2000). This theory is based on (1) rapid heating and energy loss of the surface due to evaporation, (2) sputtering and deceleration of meteoroid velocity (3) energy loss due to fragmentation and decrement of meteoroid mass and (4) internal heat conduction, which maintains uniform temperature within the meteoroids (Zinn et al. 2011).

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10 Models of the Martian Ionosphere

Meteoroid entering into the Martian atmosphere with a velocity V and angel θ is described by the equation of motion (Haider et al. 2013, Pandya and Haider 2014, Pesnall and Grabosky 2000, Molina-Cuberos et al. 2003) as given below: dz = −V · cos θ V→ = dt

(10.30)

Let m be the mass of meteoroid and δ be the effective density of the meteoroid with shape factor A then the cross section area of the effective surface of meteoroid is given by (Romig 1965; Hughes 1992): σb = A

( m )2/ 3 δ

(10.31)

The rate of change of the mass density of air during time interval dt is producing drag on the meteoroid as given below: ( m )2/ dm a 3 = σb · ρ · V = A · ρ·V dt δ

(10.32)

where ma is the mass density of the air, ρ is the atmospheric density. The meteoroids thus drag in the atmosphere and decelerate to produce rate of change of momentum as ( m )2/ dm a dp 3 =⎡·V · =⎡· A· ρ · V2 dt dt δ

(10.33)

where ⎡ is drag coefficient (Hughes 1992; Wither 2014). By the law of conservation of momentum the effective momentum of the drag produced by the air is balanced by the loss of the momentum of meteoroids as given below. −m

( m )2/ dp dV 3 = =⎡· A· ρ · V2 dt dt δ

(10.34)

Using Eq. (10.30) in Eq. (10.34) with considering the change in the velocity with height z the equation of motion can be represented as cos θ ·

⎡· A·ρ·V dV = 1 2 dz m /3 · δ /3

(10.35)

10.8 Meteoroid Ablation Model

81

10.8.1 Energy and Momentum Conservation The change in the surface temperature of meteoroid dT in the time interval dt is the result of the change in the energy dE per energy evaporation Q = c m in the same interval of time dt, where c is the heat capacity and m is the mass of meteoroid. This change in the energy can be derived in three parts: (1) The meteoroid interact with the atmospheric particles of mass ma with kinetic energy ½ ma V2 and transfer the energy to the meteoroid with heat transfer coefficient Ʌ. (2) Energy lost from the meteoroid by thermal radiation mechanism is given by ) ( 4σ ∈ T 4 − Teq4 σb (3) Energy lost due to evaporation and sputtering that causes the ablation of meteoroid mass dm in the time interval dt. Therefore, from the law of conservation of the energy, the energy transferred to the meteoroids by atmospheric particles is equal to the sum of energy lost due to meteoroid radiation, mass ablation and change in the surface temperature due to heating of meteoroid (Lebedinets et al. 1973; Rogers et al. 2005; Campball-Brown and koschny 2004) as given below: ) Q dT ( Q dm 1 Ʌρ V 3 = 4· ∈ σ T 4 − Teq4 + − 2 σb dt σb dt

(10.36)

where Ʌ is a heat transfer coefficient, ∈ is emissivity of meteoroid, σ is StefanBoltzmann constant, T is Meteoroid temperature, Teq is equivalent atmospheric temperature, Q is energy evaporation of a meteoroid (Q = cm). In the right hand side first term represents energy lost due to meteoroid radiation, second term represents change in temperature due to meteoroid heating, and third term describes mass ablation due to heating of meteoroid. The mass of meteoroid is decreasing with ablais negative. Using Eq. (10.36) we can obtain tion energy. Therefore, value of dm dt heating rate of meteoroids as given below: ( )2 ( ) 2 ) Ʌ ρ V 3 A ( m ) /3 4 ε σ A m /3( 4 dT Q dm = − (10.37) T − Teq4 + dt 2cm δ cm δ c m dt

10.8.2 Sputtering Mass Loss Levin (1946) used cathode sputtering theory to describe sputtering of a meteoroid. For small size of meteoroids thermal energy allows to replace heat transfer coefficient

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Ʌ to conveniently Ʌs as sputtering heat transfer coefficient. Thus | dm || Ʌs Q = − σb ρ V 3 dt |sp 2

(10.38)

10.8.3 Thermal Heating Mass Loss The mass loss due to sputtering process has neglected thermal ablation. In this process it is assumed that the ablation begins as soon as the surface of meteoroid starts to heat. The heat flux ablating the surface area per unit time is proportional to the atmospheric density with which the meteoroids interact. The exponential increase of the atmospheric density also increases the surface ablation exponentially, so that the surface of meteoroids observes strong evaporation with decreasing altitudes (Adolfsonet et al. 1996; Rogers et al. 2005; Bronshten 1983). The thermal heating mass loss is defined as the product of the evaporation per unit area and the area of spherical meteoroids as given below: | dm || = − εs 4π r 2 dt |ev

(10.39)

where εs is evaporation per unit area and r is a radius of spherical meteoroids. We replace π r 2 by cross sectional area σ b of uneven shape of meteoroids. Thus | dm || = − εs 4 σb dt |ev

(10.40)

The evaporation per unit area depends upon the vapour pressure Pv (T) at the evaporation temperature of the meteoroid surface. Using Langmuir expression of evaporation we have calculated εs as (Bronshten 1983; Kikwaya et al. 2011): √ εs =

μ · Pv (T ) 2π R T

(10.41)

where R is a gas constant, Pv(T) is vapour pressure and μ is molecular weight. The vapour pressure Pv (T) in above equation is calculated using Clausius-Clapeyron relation given by (Bronshten 1983; Kikwaya et al. 2011): ] [ K2 Pv (T ) = exp C 1 − T Now substitute Eq. (10.42) into (10.41) then

(10.42)

10.8 Meteoroid Ablation Model

83

[ ] μ K2 · exp C1 − 2π R T T (−k2/T ) e εs = K 1 √ T √

εs = or

(10.43)

where K1 and K2 are two parameters which define the dependence of evaporation rate at temperature T (Rogers et al. 2005; Campball-Brown and koschny 2004). The values of K1 and K2 are given below: √ K1 =

μ · eC1 2π R 2L R

and K 2 =

where L is latent heat. Using Eq. (10.43) into (10.40) then

or

| dm || e− k2 /T 4 σb = − K √ 1 dt |ev T | [ ] e− k2 /T m 2/ 3 dm || = − 4 A K √ 1 dt | δ 2/ 3 T

(10.44)

ev

The total mass loss thus obtained by sum of the mass loss due to sputtering and due to evaporation such that | | dm || dm dm || = + dt dt |ev dt | Sp (

2 m /3

− K 2/ T

dm K1e = −4 · A · √ 2 dt T δ /3 or

(

2 m /3

(10.45)

)

2 Ʌs m / 3 − A ρV 3 2/ 2Q δ 3

− K 2/ T

dm dz K1e = −4 · A · √ 2/ dz dt T δ 3

)

2 Ʌs m / 3 A − ρ V3 2/ 2Q δ 3

(10.46)

= −V · cos θ by converting vector V→ into scalar quantity V from Put V→ = dz dt Eq. (10.46) then (

cos θ

2 m /3

− K 2/ T

dm K1e = −4 · A · √ 2/ dz T 3 δ ·V

)

−

2 Ʌs m / 3 A ρV 2 2/ 2Q δ 3

(10.47)

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10 Models of the Martian Ionosphere

Using Eq. (10.46) into (10.37) we get ( )2 ) Ʌ ρ V 3 A m / 3 4 σ A ( m )2/ 3 ( 4 dT = T − Teq4 − dt 2cm δ cm δ ( ) K − 2/ T 2 ( m )2/ Ʌs m /3 K1e 3 Q− A ρ V3 −4· A · √ 2/ 2 · c · m δ T c · mδ 3

(10.48)

Using (T4 − T4 eq ) ~ T4 so that (

− K 2/ T

)

ρV A 4 A · Q · K1e dT dz 4 σ A·T = − (Ʌ − Ʌs ) − 1 2 1 2 1 2 √ / / / / dz dt 2 c m 3δ 3 c m 3δ 3 c · m /3δ /3 · T ρ V2A dT 4 σ A · T4 = cos θ (Ʌ − Ʌs ) − 1 2 dz 1 2/ c m /3δ /3 · V 2 c m / 3 δm 3 ( ) − K 2/ T 4 A · ·K 1 · Q − (10.49) √ ·e 1 2 c · m /3 · δ /3 · V T

∴

3

4

The Eqs. (10.35), (10.47) and (10.49) are used simultaneously to calculate the meteoroid impact production rates of metals in the night time and daytime ionosphere of Mars (Pandya and Haider 2012, 2014; Haider et al. 2013).

References Adolfsson, L., Gustafson, B.A.S., Murray, C.D.: The Martian atmosphere as a meteoroid detector. Icarus 119, 144–152 (1996) Banks, P.M., Nagy, A.F.: Concerning the influence of elastic scattering upon photoelectron transport and escape. J. Geophys. Res. 75(10), 1902–1910 (1970) Banks, P.M., Chappell, C.R., Nagy, A.F.: A new model for the interaction of auroral electrons with the atmosphere: spectral degradation, backscatter, optical emission, and ionization. J. Geophys. Res. 79(10), 1459–1470 (1974) Bronshten, V.A.: Physics of meteoric phenomena, D Reidel Pub Co, Holand (1983) Campbell-Brown, M.D., Koschny, D.: Model of the ablation of faint meteors. Astron. Astrophys. 418, 751–758 (2004). https://doi.org/10.1051/0004-6361:20041001 Ceplecha, Z., Borovicka, J., Elford, W.G., et al.: Meteor phenomena and bodies. Space Sci. Rev. 84, 327–471 (1998) Chen, R.H., Cravens, T.E., Nagy, A.F.: The Martian ionosphere in light of the Viking observations. J. Geophys. Res. Space. Phys. 83(A8), 3871–3876 (1978) Dalgarno, A., McElroy, M.B., Moffett, R.J.: Electron temperatures in the ionosphere. Planet Space Sci. 11(5), 463–484 (1963) Evans, S.: Inverted photoelectron diffraction: a new technique for surface structure determination. J. Electron Spectrosc. Relat. Phenom. 70(3), 217–223 (1995)

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Fisher, A.A., Hawkes, R.L., Murray, I.S., et al.: Are meteoroids really dustballs? Planet Space Sci. 48, 911–920 (2000) Fox, J.L., Yeager, K.E.: MGS electron density profiles: analysis of the peak magnitudes. Icarus 200(2), 468–479 (2009) Fox, J.L., Brannon, J.F., Porter, H.S.: Upper limits to the nightside ionosphere of Mars. Geophys. Res. Lett. 20(13), 1339–1342 (1993) Fox, J.L., Yeager, K.E.: Morphology of the near-terminator Martian ionosphere: a comparison of models and data. J. Geophys. Res. Space Phys. 111(A10) (2006) Fox, J.L.: Morphology of the dayside ionosphere of Mars: implications for ion outflows. J. Geophys. Res. Planets 114(E12) (2009) Gan, L., Cravens, T.E., Horanyi, M.: Electrons in the ionopause boundary layer of Venus. J. Geophys. Res. Space Phys. 95(A11), 19023–19035 (1990) Green, A.E.S., Jackman, C.H., Garvey, R.H.: Electron impact on atmospheric gases: 2 Yield spectra. J. Geophys. Res. 82, 5104 (1977) Haider, S.A.: Some molecular nitrogen emission from Titan-solar EUV interaction. J. Geophys. Res. Space Phys. 91(A8), 8998–9000 (1986) Haider, S.A., Mahajan, K.K.: Lower and upper ionosphere of Mars. Space Sci. Rev. 182(1), 19–84 (2014) Haider, S.A., Masoom, J.: Modeling of diffuse aurora due to precipitation of H+ -H and SEP electrons in the nighttime atmosphere of mars: monte carlo simulation and MAVEN observation. J. Geophys. Res. Space Phys. 124(11), 9566–9576 (2019) Haider, S.A., Singhal, R.P.: Analytical yield spectrum approach to electron energy degradation in Earth’s atmosphere. J. Geophys. Res. Space Phys. 88(A9), 7185–7189 (1983) Haider, S.A., Pandya, B.M., Molina-Cuberos, G.J.: Nighttime ionosphere caused by meteoroid ablation and solar wind electron-proton-hydrogen impact on Mars: MEX observation and modeling. J. Geophys. Res. Space Phys. 118(10), 6786–6794 (2013) Haider. S.A., Kim, J., Nagy, A.F., et al.: Calculated ionization rates, ion densities, and airglow emission rates due to precipitating electrons in the nightside ionosphere of Mars. J. Geophys. Res. Space Phys. 97(A7), 10637–10641 (1992) Haider, S.A., Singh, V., Choksi, V.R., Maguire, W.C., Verigin, M.I.: Calculated densities of H3 O+ (H2 O)n, NO2 − (H2 O)n, CO3 − (H2 O) n and electron in the nighttime ionosphere of Mars: impact of solar wind electron and galactic cosmic rays. J. Geophys. Res. Space Phys. 112(A12) (2007) Haider, S.A., Oyama, K.I.: Calculated electron flux and densities at 10–1000 eV in the dayside Martian ionosphere: comparison with MGS and Viking results (2002) Pandya, B.M., Haider, S.A.: Meteor impact perturbation in the lower ionosphere of Mars: MGS observations. Planet Space Sci. 63, 105–109 (2012) Haider, S.A., Bhardwaj, A.: Radial distribution of production rates, loss rates and densities corresponding to ion masses< 40 amu in the inner coma of Comet Halley: Composition and chemistry. Icarus 177(1), 196-216 (2005) Haider, S.A., Seth, S.P., Brain, D.A., et al.: Modeling photoelectron transport in the Martian ionosphere at Olympus Mons and Syrtis Major: MGS observations. J. Geophys. Res. Space Phys. 115(A8), (2010) Haider, S.A., Mahajan, K.K., Kallio, E.: Mars ionosphere: a review of experimental results and modeling studies. Rev. Geophys. 49(4) (2011) Hanson, W.B., Patterson, T.N.L., Degaonkar, S.S.: Some deductions from a measurement of the hydrogen ion distribution in the high atmosphere. J. Geophys. Res. 68(22), 6203–6205 (1963) Hanson, W.B., Johnson, F.S.: Electron temperatures in the ionosphere. In Liege International Astrophysical Colloquia, vol. 10, pp. 390–424 (1961) Hughes, D.W.: The meteor flux. Space Sci Rev 61, 275–299 (1992) Hughes, D.W.: Meteors. Cosmic dust, pp 123–185 (1978) Ip, W.H.: On a hot oxygen corona of Mars. Icarus 76(1), 135–145 (1988) Jackman, C.H., Green, A.E.S.: Electron impact on atmospheric gases 3 Spatial yield spectra for N2 . J. Geophys. Res. 84, 2715–2724 (1979)

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Kallio, E., Liu, K., Jarvinen, R., et al.: Oxygen ion escape at Mars in a hybrid model: high energy and low energy ions. Icarus 206(1), 152–163 (2010) Kallio, E., Barabash, S.: Atmospheric effects of precipitating energetic hydrogen atoms on the Martian atmosphere. J. Geophys. Res. Space Phys. 106(A1), 165–177 (2001) Kallio, E., Janhunen, P.: Ion escape from Mars in a quasi-neutral hybrid model. J. Geophys. Res. Space Phys. 107(A3), SIA-1 (2002) Kikwaya, J.-B., Campbell-Brown, M., Brown, P.G.: Bulk density of small meteoroids. Astro. Astrophy. 530, A113 (2011). https://doi.org/10.1051/0004-6361/201116431 Lebedinets, V.N., Manochina, A.V., Shushkova, V.B.: Interaction of the lower thermosphere with the solid component of the interplanetary medium. Planet Space Sci. 21, 1317–1332 (1973) Levin, B.Yu.: Velocities, orbits, and masses of meteorites. Astron, Zhurn 23, 83–96 (1946) Luhmann, J.G.: Auroral bremsstrahlung spectra in the atmosphere. J. Atmos. Terr. Phys. 39(5), 595–600 (1977) Ma, Y., Nagy, A.F., Sokolov, I.V., Hansen, K.C.: Three-dimensional, multispecies, high spatial resolution MHD studies of the solar wind interaction with Mars. J. Geophys. Res. Space Phys. 109(A7), (2004) Mantas, G.P., Hanson, W.B.: Photoelectron fluxes in the Martian ionosphere. J. Geophys. Res. Space Phys. 84(A2), 369–385 (1979) Mantas, G.P., Hanson, W.B.: Evidence of solar wind energy deposition into the ionosphere of Mars. J. Geophys. Res. Space Phys. 90(A12), 12057–12064 (1985) Mantas, G.P., Hanson, W.B.: Analysis of Martian ionosphere and solar wind electron gas data from the planar retarding potential analyzer on the Viking spacecraft. J. Geophys. Res. Space Phys. 92(A8), 8559–8569 (1987) Ma, Y., et al.: Three-dimensional, multispecies, high spatial resolution MHD studies of the solar wind interaction with Mars. J. Geophys. Res. Space Phys. 109(A7) (2004) Molina-Cuberos, G.J., Witass, O., Lebreton, J.P., et al.: Meteoric ions in the atmosphere of Mars. Planet Space Sci. 51(3), 239–249 (2003) McKinley, D.W.R.: Meteor Science and Engineering, McGraw-Hill Book Co., New York, p. 309, (1961) Nagy, A., Ma, Y., Sokolov, I.: 3D, multi-species, high spatial resolution MHD studies of the solar wind interaction with Mars. 35th COSPAR Scientific Assembly, 35, (2004) Öpik, E.J.: Physics of Meteor Flight in the Atmosphere, Interscience Publishers Inc., New York, p. 174, (1958) Poter, H.S.: Analytic total and angular elastic electron cross sections for planetary, atmospheres. Computer Science Corp. Report CSC/TM-78/6017 (1978). Pandya, B.M., Haider, S.A.: Numerical simulation of the effects of meteoroid ablation and solar EUV/X-ray radiation in the dayside ionosphere of Mars: MGS/MEX observations. J. Geophys. Res. Space. Phys. 119(11), 9228–9245 (2014) Pesnell, W.D., Grebowsky, J.: Meteoric magnesium ions in the Martian atmosphere. J. Geophys. Res. 105, 1695–1707 (2000) Riewe, F., Green, A.E.: Ultraviolet aureole around a source at a finite distance. Appl. Opt. 17(12), 1923–1929 (1978) Rogers, L.A., Hill, K.A., Hawkes, R.L.: Mass loss due to sputtering and thermal processes in meteoroid ablation. Planet Space Sci. 53(13), 1341–1354 (2005) Rohrbaugh, R.P., Nisbet, J.S., Bleuler, E., Herman, J.R.: The effect of energetically produced O2+ on the ion temperatures of the Martian thermosphere. J. Geophys. Res. Space Phys. 84(A7), 3327–3338 (1979) Romig, M.F.: The physics of Meteor entry. AIAA Journal 3(3), 385–394 (1965) Shah, S.Y., et al.: A coupled model of the D and E regions of Mars’ ionosphere for flare and non-flare electron density profiles. Icarus 361, 114403 (2021) Shinagawa, H.: A two-dimensional model of the Venus ionosphere: 2 Magnetized ionosphere. J. Geophys. Res. Space Phys. 101(A12), 26921–26930 (1996)

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Shinagawa, H., Bougher, S.W.: A two-dimensional MHD model of the solar wind interaction with Mars. Earth, Planets Space 51(1), 55–60 (1999) Shinagawa, H., Cravens, T.E.: A one-dimensional multispecies magnetohydrodynamic model of the dayside ionosphere of Mars. J. Geophys. Res. Space Phys. 94(A6), 6506–6516 (1989) Singhal, R.P., Green, A.E.S.: Spatial aspects of electron energy degradation in atomic oxygen. J. Geophys. Res. Space Phys. 86(A6), 4776-4780 (1981) Singhal, R.P., Haider, S.A.: Some molecular nitrogen emissions from Titan-Solar EUV & magnetospheric interaction (1986) Singhal, R.P., Haider, S.A.: Analytical yield spectrum approach to photoelectron fluxes in the Earth’s atmosphere. J. Geophys. Res. Space. Phys. 89(A8), 6847–6852 (1984) Singhal, R.P., Jackman, C.H., Green, A.E.S.: Spatial aspects of low-and medium-energy electron degradation in N2 . J. Geophys. Res. Space Phys. 85(A3), 1246–1254 (1980) Stamnes, K., Swanson, R.A.: A new look at the discrete ordinate method for radiative transfer calculations in anisotropically scattering atmospheres. J. Atmos. Sci. 38(2), 387–399 (1981) Thirupathaiah, P., Shah, S.Y., Haider, S.A.: Characteristics of solar X-ray flares and their effects on the ionosphere and human exploration to Mars: MGS radio science observations. Icarus 330, 60–74 (2019) Whipple, F.L.: A comet model I The acceleration of Comet Encke. Astrophys. J. 111, 375, ISSN: 0004-637X, 1538-4357. https://doi.org/10.1086/145272 Whipple, F.L.: A comet model II physical relations for comets and meteors. Astrophys. J. 113, 464. https://doi.org/10.1086/145416 Willmore, A.P.: Ionospheric heating in the F-region. Proc. r. Soc. Lond. A 281(1384), 140–149 (1964) Wither P (2014) Predictions of the effects of Mars’s encounter with cometC/2013A1 (Siding Spring) upon metalspecies in its ionosphere. Geophys. Res. Lett. 41 (2014). https://doi.org/10.1002/201 4GL061481 Zinn, J., Close, S. Colestock, P.L., MacDonell, A., Loveland, R.: Analysis of ALTAIR 1998 meteor radar data. J. Geophys. Res. 116, A04312 (2011). https://doi.org/10.1029/2010JA015838

Chapter 11

Solar Flux for Ionospheric Modeling of Mars

Abstract The solar flux is the basic indicator to measure solar radiation. The solar radiation received at the top of the Martian atmosphere is absorbed by the atmospheric species depending upon the component rays of sunlight and their wavelengths. The flux at any point in the atmosphere depends on the properties and quantity of absorbing gases at higher altitudes. There are several measurements and models for the solar EUV and X-ray fluxes. In this chapter we have discussed the measurements and models of solar EUV and X-ray fluxes carried out by SOHO, SORCE, GOES, SOLAR 2000 and FISM. Keywords EUV flux · X-ray flux · Solar spectra

11.1 EUV Flux Measurements: SOHO SOHO is a European Space Agency (ESA) spacecraft (Judge et al. 1998). The EUV and XUV photon fluxes are measured by the Solar EUV Monitor (SEM) experiment onboard SOHO (Meier et al. 2002; Tsurutani et al. 2005; Liu et al. 2007). It measures EUV flux integrated in the wavelength band 26 to 34 nm and the XUV flux integrated in the wavelength band 0.1 to 50 nm. The Fig. 11.1 represents the solar EUV and soft X-ray fluxes measured by SOHO in the wavelength 0.1–50 nm. Time intervals of the operation of MARSIS, IMA/ASPERA-3 and MAVEN are also shown in this figure. The solar flux varies approximately from 0.6 × 1010 photons/cm2 s to 2 × 1010 photons/cm2 s Dubinin et al. 2017). The SOHO data is freely accessible on ftp://pds-geosciences.wustl.edu/dep/space-science/semdata.html. The SEM-SOHO fluxes have been also used to study the effect of solar flare and CMEs in Martian ionosphere (Withers and Mendillo 2005; Liu et al. 2007; Futaana et al. 2008; Mahajan et al. 2009; Haider et al. 2009; Ulusen et al. 2012; Dubinin et al. 2016; Lee et al. 2018; Li et al. 2022).

© Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_11

89

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11 Solar Flux for Ionospheric Modeling of Mars

Fig. 11.1 Solar EUV flux (0–50 nm) measured by SEM/SOHO at the Earth orbit and translated to Mars position during the MEX mission. Black and green lines correspond to the mean and median values, respectively (from Dubinin et al. 2017)

11.2 X-ray Flux Measurements: GOES The Geostationary Operational Environmental Satellite (GOES) is equipped with two X-ray sensors, one for monitoring the 0.5 to 4 Å, or short band, and one for the 1 to 8 Å, or long band which measures the solar X-ray flux (Bornmann et al. 1996). The data are collected at two-second intervals and compiled by NOAA’s Space Weather Prediction Centre (SWPC) into one minute and five minute averages for public use. The data are grouped by day from 00:00:00 to 23:59:59 UT. The Fig. 11.2 represents the GOES spectra for the 0.05–0.3 nm and 0.1 − 0.8 nm wavelengths. These fluxes are varying from 10–6 to 10–8 W/m2 for both wavelengths. The GOES X-ray flux data is freely accessible on https://www.ngdc.noaa.gov/stp/satellite/goes/dataaccess. html. Since the 1970s, the GOES series of satellites have been continuously observing the solar X-ray fluxes. GOES data at 0.1–0.8 nm are also useful for identifying solar

11.3 EUV and X-ray Flux Measurements: SORCE

91

Fig. 11.2 Solar X-ray flux in two bands (0.05–0.3 nm and 0.1 − 0.8 nm) observed by GOES-11 satellite. (https://www.ngdc.noaa.gov/ stp/satellite/goes/dataaccess. html)

flares, but these wavelengths are shorter than those responsible for most ionization at higher altitude. GOES X-ray flux can be enhanced by orders of magnitude during a flare time. Several investigators have modelled the solar x-ray flares on Mars by using the GOES X-ray flux (for e.g., Mendillo et al. 2006; Haider et al. 2009, 2012; Fellows et al. 2015; Haider et al. 2016; Thirupathaih et al. 2019; Shah et al. 2021).

11.3 EUV and X-ray Flux Measurements: SORCE The Solar Radiation and Climate Experiment (SORCE) observed solar irradiance with for ultraviolet, visible and near-infrared wavelength range 1–40 nm, 115– 310 nm, and 310–2400 nm. SORCE was a NASA mission that operated from 2003 to 2020 to provide key climate-monitoring measurements of Total Solar Irradiance (TSI) and Solar Spectral Irradiance (SSI) (Rottman et al., 2006). The SORCE database is publicly accessible at http://lasp.colorado.edu/lisird/sorce/ The SORCE/SSI instruments include the Spectral Irradiance Monitor (SIM) for 240–2413 nm range, SOLar STellar Irradiance Comparison Experiment (SOLSTICE) for the 115–308 nm range, and X-ray UV Photometer System (XPS) for the 0.1–40 nm range and for the H I 121.6 nm emission. The Fig. 11.3 shows the 11 years variability of TSI from 2003 to 2014 observed by the SORCE (Bonal et al. 2016; Woods et al. 2021). The source data has also been used to calculate the ion density of meteoric metals and electron density in atmosphere of Mars (Pandya and Haider 2014).

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11 Solar Flux for Ionospheric Modeling of Mars

Fig. 11.3 Total Solar Irradiance (TSI) data during 2003 to 2014 from SORCE (from Bonal et al. 2016)

11.4 EUV and X-ray Model Flux: SOLAR 2000 Tobiska et al. (2000) developed the SOLAR2000 model (S2K) based on the measurements of the solar irradiance provided by several satellites and rockets from 1976 to 1998. The S2K is an empirical solar irradiance specification tool for accurately characterizing solar irradiance variability across the solar spectrum. It provides solar irradiance for any wavelength between 1 and 106 nm for any given day. The S2K is designed to be a fundamental energy input into planetary atmosphere models and a tool to model or predict the solar radiation component of the space environment. Tobiska et al. (2008) also developed the Solar Irradiance Platform (SIP) which incorporates the SOLAR2000 model and other models. The SIP (formerly known as Solar2000 model) is an empirical solar irradiance specification tool for accurately characterizing solar irradiance variability across the solar spectrum. The SIP is designed to be a fundamental energy input into planetary atmosphere models and a tool to model or predict the solar radiation component of the space environment. The S2K has been well tested in the modeling of dayglow emissions (Thirupathaiah et al. 2022). Previous studies showed that in the absence of solar X-ray flux measurements, SOLAR 2000 model was used to calculate the production rate, ion and electron densities in the ionosphere of Mars (Martinis et al 2003; Fox and Yeager 2006; Lolo et al. 2012; Peter et al. 2014, 2021). This model predicts the solar irradiance variability across the soft X-ray, EUV spectrum through Lyman alpha (121.6 nm), visible and IR which is given in Fig. 11.4 at quiet condition.

11.5 FISM Model Flux for X-ray Flares

93

Fig. 11.4 Solar irradiance for SOLAR2000 model from the X-rays to infrared radiation (from Tobiska 2000)

11.5 FISM Model Flux for X-ray Flares FISM is a soft X-ray Flare Irradiance Spectral Model (FISM), which estimates X-ray flux at 5–90 Å (Chamberlin et al. 2006). This model did not estimate X-ray flux at 4 Å. FISM is more appropriate for flares because of its high spectral and temporal resolution in comparison to other solar irradiance nodels (Hinterreger et al. 1981; Richards et al. 2006; Tobiska et al. 2000). Recently Shah et al. (2021) used combined fluxes of GOES and FISM for wavelength range 0.5–90 Å. The Figs. 11.5a and 11.5b represent the spectra of GOES (0.5–3 Å) and FISM (5–90 Å) on 6 April, 2001 and 17 March, 2003. The strong solar X-ray flares occurred on both days. FISM represents peak X-ray flux at 15 Å. GOES X-ray peak flux is nearly same at 0.5 Å and 1–8 Å as estimated by FISM (GOES flux for wavelength band 1–8 Å is not plotted in these figures). GOES did not measure X-ray flux at higher wavelengths beyond 8 Å (Haider and Mahajan 2014). GOES X-ray peak flux at 0.5–3 Å is lower by two orders of magnitude than that estimated by FISM X-ray peak flux. The X-ray flux is not changing much with UT between pre- and post flare periods. Shah et al. (2021) estimated the electron density profiles using GOES and FISM fluxes for flare and non-flare period.

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Fig. 11.5 (a) the time series of GOES and FISM spectrum plotted by different colours at different wavelengths on 6 April 2001, (b) same spectrum as shown in Fig. 11.5a but for 17 March 2003 (from Shah et al. 2021)

References Bornmann, P.L., Speich, D., Hirman, J., et al., GOES solar X-ray imager: overview and operational goals. GOES-8 and Beyond 2812, 309–319 (1996) Chamberlin, P.C., Woods, T.N., Eparvier, F.G.: Flare Irradiance Spectral Model (FISM) use for space weather applications. In: Proceedings of the ILWS Workshop, vol. 153 (2006) Delgado-Bonal, A., Zorzano, M.P., Martín-Torres, F.J.: Martian top of the atmosphere 10–420 nm spectral irradiance database and forecast for solar cycle 24. Sol. Energy 134, 228–235 (2016) Dubinin, E., Fraenz, M., Andrews, D., Morgan, D.: Martian ionosphere observed by Mars Express 1 influence of the crustal magnetic fields. Planet Space Sci. 124, 62–75 (2016) Dubinin, E., Fraenz, M., Pätzold, M., et al.: Martian ionosphere observed by Mars Express 2 influence of solar irradiance on upper ionosphere and escape fluxes. Planet Space Sci. 145, 1–8 (2017) Fallows, K., Withers, P., Gonzalez, G.: Response of the Mars ionosphere to solar flares: analysis of MGS radio occultation data. J. Geophys. Res. Space Phys. 120(11), 9805–9825 (2015) Fox, J.L., Yeager, K.E.: Morphology of the near-terminator Martian ionosphere: a comparison of models and data. J. Geophys. Res. Space Phys. 111(A10) (2006) Futaana, Y., Barabash, S., Yamauchi, M., et al.: Mars Express and Venus Express multi-point observations of geoeffective solar flare events in December 2006. Planet Space Sci. 56(6), 873–880 (2008) Haider, S.A., Mahajan, K.K.: Lower and upper ionosphere of Mars. Space Sci. Rev. 182(1), 19–84 (2014) Haider, S.A., Batista, I.S., Abdu, M.A., et al.: Flare X-ray photochemistry of the E region ionosphere of Mars. J. Geophys. Res. Space Phys. 121(7), 6870–6888 (2016)

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Haider, S.A., Abdu, M.A., Batista, I.S., et al.: On the responses to solar X-ray flare and coronal mass ejection in the ionospheres of Mars and Earth. Geophys. Res. Lett. 36(13) (2009) Haider, S.A., McKenna-Lawlor, S.M.P., Fry, C.D., Jain, R., Joshipura, K.N.: Effects of solar X-ray flares in the E region ionosphere of Mars: first model results. J. Geophys. Res. Space Phys. 117(A5) (2012) Hinteregger, H.E., Fukui, K., Gilson, B.R.: Observational, reference and model data on solar EUV, from measurements on AE-E. Geophys. Res. Lett. 8(11), 1147–1150 (1981) Judge, D.L., McMullin, D.R., Ogawa, H.S., et al.: First solar EUV irradiances obtained from SOHO by the CELIAS/SEM. In: Solar Electromagnetic Radiation Study for Solar Cycle 22, pp. 161– 173. Springer, Dordrecht (1998) Lee, C.O., Jakosky, B.M., Luhmann, J.G., et al.: Observations and impacts of the 10 September 2017 solar events at Mars: an overview and synthesis of the initial results. Geophys. Res. Lett. 45(17), 8871–8885 (2018) Li, X., Wang, Y., Guo, J., Lyu, S.: Solar energetic particles produced during two fast coronal mass ejections. Astrophys. J. Lett. 928(1), L6 (2022) Liu, L., Wan, W., Yue, X., et al.: The dependence of plasma density in the topside ionosphere on the solar activity level. Annales Geophysicae 25(6), 1337–1343 (2007). Copernicus GmbH Lollo, A., Withers, P., Fallows, K., et al.: Numerical simulations of the ionosphere of Mars during a solar flare. J. Geophys. Res. Space Phys. 117(A5) (2012) Mahajan, K.K., Lodhi, N.K., Singh, S.: Ionospheric effects of solar flares at Mars. Geophys. Res. Lett. 36(15) (2009) Martinis, C.R., Wilson, J.K., Mendillo, M.J. Modeling day-to-day ionospheric variability on Mars. J. Geophys. Res. Space Phys. 108(A10) (2003) Meier, R.R., Warren, H.P., Nicholas, A.C., et al.: Ionospheric and dayglow responses to the radiative phase of the Bastille Day flare. Geophys. Res. Lett. 29(10), 99–101 (2002) Mendillo, M., Withers, P., Hinson, D., et al.: Effects of solar flares on the ionosphere of Mars. Science 311(5764), 1135–1138 (2006) Pandya, B.M., Haider, S.A.: Numerical simulation of the effects of meteoroid ablation and solar EUV/X-ray radiation in the dayside ionosphere of Mars: MGS/MEX observations. J. Geophys. Res. Space Phys. 119(11), 9228–9245 (2014) Peter, K., Pätzold, M., Molina-Cuberos, G., et al.: The dayside ionospheres of Mars and Venus: comparing a one-dimensional photochemical model with MaRS (Mars Express) and VeRa (Venus Express) observations. Icarus 233, 66–82 (2014) Peter, K., Pätzold, M., Molina-Cuberos, G.J.: The lower dayside ionosphere of Mars from 14 years of MaRS radio science observations. Icarus 359, 114213 (2021) Richards, P.G., Woods, T.N., Peterson, W.K.: HEUVAC: a new high resolution solar EUV proxy model. Adv. Space Res. 37(2), 315–322 (2006) Shah, S.Y., et al.: A coupled model of the D and E regions of Mars’ ionosphere for flare and non-flare electron density profiles. Icarus 361, 114403 (2021) Thirupathaiah, P., Shah, S.Y., Haider, S.A.: Characteristics of solar X-ray flares and their effects on the ionosphere and human exploration to Mars: MGS radio science observations. Icarus 330, 60–74 (2019) Thirupathaiah, P., Haider, S.A., Masoom, J.: Simulation of photoelectron flux, electron density, emission rate and limb intensity of CO (a3π) Cameron bands in the Martian thermosphere: Comparisons with (1) SPICAM and IUVS observations and (2) other model calculations. J. Earth Syst. Sci. 131(2), 1–15 (2022) Tobiska, W.K., Bouwer, S.D., Bowman, B.R.: The development of new solar indices for use in thermospheric density modeling. J. Atmos. Solar Terr. Phys. 70(5), 803–819 (2008) Tobiska, W.K., Woods, T., Eparvier, F., et al.: The SOLAR 2000 empirical solar irradiance model and forecast tool. J. Atmos. Sol. Terr. Phys. 62, 1233–1250 (2000) Tsurutani, B.T., Judge, D.L., Guarnieri, F.L., et al.: The October 28, 2003 extreme EUV solar flare and resultant extreme ionospheric effects: comparison to other Halloween events and the Bastille Day event. Geophys. Res. Lett. 32(3) (2005)

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Ulusen, D., Brain, D.A., Luhmann, J.G., Mitchell, D.L.: Investigation of Mars’ ionospheric response to solar energetic particle events. J. Geophys. Res. Space Phys. 117(A12) (2012) Withers, P., Mendillo, M.: Response of peak electron densities in the Martian ionosphere to dayto-day changes in solar flux due to solar rotation Planet. Space Sci. 53(14–15), 1401–1418 (2005)

Chapter 12

Ionization Sources of Upper Ionosphere of Mars

Abstract Since photoionization serves as a dominant source of ionization, it is not surprising that the structure of the dayside Martian ionosphere is strongly modulated by the incident solar illumination condition in terms of solar EUV and X-ray irradiance. UV and X-ray radiations are the major ionizing sources in the upper atmosphere of Mars. The upper ionosphere of Mars is divided into E and F region. The E and F regions are produced in the upper ionosphere at about 110 and 130 km due to impact of soft X-rays (10–100 Å) and solar EUV radiation (100–1026 Å) respectively. The ionization depends primarily on solar radiation and suns’ activity. There is a diurnal and seasonal variability effects on Mars’ atmosphere. Mars is away and close to the sun during northern winter and northern summer season. Thus it will receive less solar UV radiation during winter than summer. The solar activity of the sun is also associated with the sunspot cycle. During more sunspots more radiation occurs from the sun. Energetic electron impact and charge-exchange processes also contribute more to the production of planetary ions during high solar activity period. Keywords Ionospheric processes · Parameters · Ionospheric layes

12.1 Ionization by Solar EUV: F Region Ionosphere The F region is formed due to photoionization of neutral atmosphere of Mars by solar EUV radiation from 90 to 1026 Å. The major gases in the upper ionosphere of Mars viz. CO2 , N2 , O2 , O, Ar and CO are ionized by solar EUV radiation. The maximum ionization in the F region occurs at altitude ∼ 125–135 km in the dayside ionosphere of Mars for sub-solar condition (Hanson et al. 1977; Fox and Dalgarno 1979; Bougher et al. 1990; Morgan et al. 2008; Fox and Yeager 2009; Fox and Weber 2012; Haider and Mahajan 2014; Fallows et al. 2015; Withers et al. 2015; Vogt et al. 2017; Yao et al. 2019; Thirupathaiah et al. 2019; Shah et al. 2021). The F peak varies with Chapman theory. At higher solar zenith angle the F peak ionization occurs at higher altitudes. According to Chapman theory, the F peak is located, where the optical depth is unity. The F peak also varies with solar cycles and solar rotation (Bougher et al. 2001, 2004; Forbes et al., 2006; Rao et al. 2014; Hughes et al. 2022). © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_12

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Alternatively, the variation could also occur on relatively short timescales, such as during XUV flare events and global dust storms (Mahajan et al. 2009; Withers and Pratt 2013; Andrews et al. 2015; Withers et al. 2018; Felici et al. 2020; Girazian et al. 2020).

12.2 Ionization by X-rays: E Region Ionosphere The E region is formed due to ionization by soft X rays characterized by wavelengths 10–90 Å at about ~ 110–115 km (Fox 2004; Haider et al. 2009a, b, 2016; Fallows et al. 2015; Thirupathaiah et al. 2019; Shah et al. 2021). At night the E layer disappears because the primary source of ionization is no longer present. Figure 12.1 shows the electron density profiles from ROSE experiment onboard MAVEN on 9 August (red curve) and 16 August (black curve), 2016 at SZA 70° and 67° respectively (Withers et al. 2018). The E peak density on 16 August, 2016 increases by a factor of 1.7 from 9 August, 2016. This is due to the fact that ionizing flux is more by a factor of 2.8 on 16 August from 9 August, 2016. The measured E peak density is 3 × 1010 m−3 on 9 August, 2016 but 5 × 1010 m−3 on 16 August 2016. The F peak density is nearly same as 1.1 × 1011 m−3 for both days.

12.3 Solar Wind Impact Ionization Mars has an induced magnetosphere outside the crustal magnetic field region. Therefore, solar wind dynamic pressure compresses the IMF field into the Martian ionosphere. A one-dimensional MHD model has been developed by Shinagawa and Cravens (1989) to study the role of electromagnetic forces in the Martian ionosphere. They found a good agreement between model and Viking observations after adding an extra heat source produced due to the solar wind in the upper ionosphere. Later Shinagawa and Bougher (1999) developed a two-dimensional MHD model and studied two cases of solar wind dynamic pressure on Mars. Ma et al. (2004) have developed a three-dimensional model. They reported that solar wind plays an important role above 250 km. Haider et al. (2010) also reported similar conclusions by using a one-dimensional model with non-zero upward flux boundary condition. First measurement of nightside ionosphere of Mars was carried out by RO experiment onboard Mars 4 and Mars 5 at SZA 127° and 106° in February 1974 during solar minimum condition (Savich and Samovol 1976). Later Zhang et al. (1990) analyzed Viking data and have reported that about 60% of the RO profiles did not show a welldefined peak during the nighttime at low solar activity condition. Remaining 40% profiles observed a peak value 5 × 103 cm−3 at an altitude of about 150 km during the nighttime ionosphere. This peak is produced due to solar wind electron transportation from dayside to nightside atmosphere across the terminator (cf. Verigin et al. 1991;

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Fig. 12.1 ROSE measurements from MAVEN on 9 August, 2016 (SZA ~ 70°, red line) and 16 August, 2016 (SZA ~ 67°, black line). Horizontal lines indicate 1-sigma error in the electron density (from Withers et al. 2018)

Haider et al. 1992; Fox et al. 1993; Haider 1997; Cui et al. 2009; González-Galindo et al. 2013; Haider and Mahajan 2014; Chaufray et al. 2014; Ma et al. 2015). The H+ –H impact is also reported an important source of the nighttime ionosphere of Mars (Kallio and Barabash 2001; Kallio and Janhunen 2001, Haider et al. 2002). Recently Haider and Masoom (2019) used this source of ionization and produced electron densities ~ 3.5 × 103 cm−3 and 2.0 × 103 cm−3 at SZA 105° and 127o respectively. They found that H+ –H penetrate deeper into Martian atmosphere and lose their energy at lower altitude as compared to solar wind electron impact ionization.

12.4 Dynamics of the Upper Ionosphere The dynamics of the upper ionosphere of Mars is controlled by EUV and soft X-ray of the solar spectrum. Therefore, the upper ionosphere of Mars is expected to vary significantly with long and short time scales. The irradiance is highly variable on short time scales from minutes to hours due to solar flares, days to weaks due to solar

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rotation and years due to 11 years solar cycle (Girazian and Withers 2015; Haider et al. 2009c, 2022a, b; Haider and Mahajan 2014). The Mars’ ionosphere is also predicted to vary with season and dust. Mars is a very windy planet. Therefore, various types of atmospheric motions are also expected to exist in the troposphere and upper atmosphere of Mars. The wind is composed of GW, tidal waves, planetary waves and stationary waves. In the Mars’ ionosphere the GW dissipate rapidly due to molecular viscosity and thermal conduction because of their short wavelength (Imamura and Ogawa 1995; Rao et al. 2021, 2022). The tidal oscillation is a short wave which produces several modes (1–7 modes) of the motion in the Mars’ atmosphere (Bougher et al. 2001; Haider et al. 2006, 2009a, b, 2010; Seth et al. 2006a, b; Cavalie et al. 2008; Haider et al. 2009c, 2010). The planetary wave is a long wave, which is produced due to rotating the planet. Figure 12.2 represents the nighttime zonal wave distribution of temperature observed by MGS at altitude ~ 30 km during summer at latitude range 64.7° N– 67.3° N (Haider et al. 2009a, b). Figure 12.3 represents the nighttime zonal wave distribution of density observed by MGS at the same altitude and latitude during the summer as shown in Fig. 12.2. The dotted and dashed lines in both figures represent lower and upper 0.95 confidence limits. The error bar is also plotted in these figures. The solid lines represent the best fit to the observation. The black circle represents the observations with error bar. These measurements suggest that the mode of wave number 2 of semi-diurnal tides is dominated in the tropospheric temperature and densities during summer at northern high latitudes (Haider et al. 2009). The troughs and sink in the longitudinal distrbution profiles of Figs. 12.2 and 12.3 represent the modes of the atmospheric tides.

12.5 Chemistry of the Upper Ionosphere Solar X-ray and EUV radiation ionizes the neutral species in the E and F region of Mars respectively. These radiations produced msainly CO2 + , O2 + , O+ , CO+ , NO+ and N2 + ions in the upper ionosphere. The chemistry of these ions has been studied by several investigators (cf. Chen et al. 1978; Hanson et al. 1977; Fox et al. 1993; Haider 1997; Ma et al. 2004; Duru et al. 2008; Haider et al. 2010, 2012, 2016; Thirupathaiah et al. 2019). The photoionization rate of CO2 + is a dominant process in the upper ionosphere of Mars. This ion is mainly destroyed by atomic oxygen and produced a dominant ion O2 + in the F region. The ion O2 + is later destroyed by dissociative recombination process. The ion NO+ is produced in the E region due to impact of O2 + with N and NO (Thirupathaiah et al. 2019; Shah et al. 2021) and it is also destroyed by dissociative recombination reaction. The major ions are O2 + , NO + and CO2 + below 200 km. Above this altitude O+ ion dominates. The density of NO+ is directly proportional to the densities of N and NO (Haider et al. 2016). The other major source of this ion is the reaction of CO+ and N2 + with NO. Among the minor ions, CO+ is lost by the charge exchange reaction with CO2 . This process

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Fig. 12.2 Zonal structures of summer temperature measured by MGS between longitudes 0° and 360° E and latitudes 64.7° N and 67.3° N at altitude 30 km. The solid line represents the best fit to the data. The dotted and dashed lines represent lower and upper 0.95 confidence limits. The error bar is also plotted (from Haider et al. 2009c)

Fig. 12.3 Zonal structures of density measured by MGS during summer longitudes 0° and 360° E and latitudes 64.7° N and 67.3° N at altitude 30 km. The solid line represents the best fit to the data. The dotted and dashed lines represent lower and upper 0.95 confidence limits. The error bar is also plotted (from Haider et al. 2009c)

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destroys almost all CO+ ions in the Martian ionospher. The chemical reactions for the upper ionospheric chemistry are given in Appendix E (Haider et al. 2016).

12.6 Modeling of the Upper Ionosphere Various physical, chemical and dynamical processes have been used in MTGCM by Bougher et al. (2000, 2015a, 2017; b) to study the upper ionosphere of Mars. This model calculates self-consistently neutral, ion, electron temperatures, and neutral/ ion densities. It also calculates three-component of neutral winds over the Martian globe under solar maximum, medium and minimum activity during different Mars’ seasons. This model also accommodates diurnal and semi-diurnal tidal mode amplitudes and phases consistent for dust storm events. The heating due to regional or global dust storms modifies the Martian tidal parameters significantly, which are greatly enhancing the tidal impact on Mars’ thermosphere during dust storms. The dust influence is very strong during southern summer solstices at Ls = 200–300 (Haider et al. 2022a, b). Ma et al. (2004) and Duru et al. (2008) have used threedimensional models to study the dynamics of the upper ionosphere of Mars. They found that solar wind plays an important role above 250 km, therefore the ion density profiles begin to differ from their chemical equilibrium values. MCD provides meteorological parameters like air density, temperature, wind and mixing ratios using GCM at different altitude, latitude and longitude in the Martian atmosphere (Millour et al. 2014). This model has been also developed for low, medium and high dust storm conditions. The photochemical model for the calculation of ion and electron densities is also developed (Shimazaki 1989; Rodrigoe et al. 1990; Haider 1997; Bhardwaj and Michael 1999; González-Galindo et al. 2005, 2013; Haider et al. 2009a, b, 2011, 2016; Mendillo et al. 2015, 2017; Mukundan et al. 2020; Shah et al. 2021; Fox et al. 2021; Wu et al. 2021). Figure 12.4 shows comparison of average of NGIMS and LPW measurements with three models: (1) MIRI (purple line), (2) Chapman (blue line) and (3) Nemec et al. (2016) (green line). The NGIMS measures the ion density and LPW observes the electron density. Under charge neutrality condition the sum of NGIMS ions and LPW electron density should be nearly same. The average of these two measurements provides a good representative of average electron density. These observations were carried out on 21 April, 2015 during inbound orbit # 1082 (17:04:03–17:23:41 UT) and outbound orbit # 1083 (17:14:03–17:23:41 UT) (Mendillo et al. 2017). These models and observations were carried out at nearly same SZA ~ 25°.

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Fig. 12.4 Average of NGIMS and LPW results shown in comparison to the three models for 21 April 2015 conditions. The average Nmax from LPW and NGIMS is 15.3 × 104 e cm−3 . The average Nmax from MIRI and Nˇemec is 18.1 × 104 e cm−3 (from Mendillo et al. 2017)

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Chapter 13

Mars Upper Ionospheric Disturbances

Abstract Mars has an induced magnetosphere, where solar wind interacts directly with the upper ionosphere/atmosphere and can induce a series of solar disturbances from solar flares, SEPs and CMEs (Haider et al. in J Geophys Res Space Phys 114(A3) 2009a, Geophys Res Lett 36(13) 2009b; Haider et al. in J Geophys Res Space Phys 117(A5), 2012; Thirupathaiah et al. in Icarus 330:60–74, 2019; Shah et al. 2021). The GW and Travelling Ionospheric Disturbances (TID) also occur in the upper ionosphere of Mars (Rao et al. 2022, 2023; Nakagawa et al. in J Geophys Res Planets 125(9):e2020JE006481, 2020; Zhang et al. in J Geophys Res Space Phys. 124:5894–5917, 2019; Collinson et al. Geophys Res Lett 46:4554–4563, 2019). A solar flare is an intense burst of radiation coming from the release of magnetic energy associated with sunspots, SEPs also emit from the sun as bursts of high energy particles lasting for hours or sometimes days. CMEs are large expulsions of plasma and magnetic field from sun’s corona. GWs are small scale disturbances in the atmospheric variables such as pressure, density, wind and temperature. TIDs are the ionospheric manifestation of atmospheric GWs in the upper atmosphere. Keywords Solar flare · CMEs · SEPs

13.1 Effects of Solar Flares on the Upper Ionosphere MGS observed 5600 electron density profiles during solar maximum conditions. Among these profiles the responses of X, M and C class solar flares were observed in the E layer at altitude 110 km from 32 electron density profiles obtained from RO experiment (Fallows et al. 2015; Thirupathaiah et al. 2019). During the flare events the peak electron density increased suddenly by a factor of 5 (Mendillo et al. 2006; Mahajan et al. 2009; Haider et al. 2016; Fallows et al. 2015; Thirupathaiah et al. 2019; Shah et al. 2021). Later MARSIS observed an increase in the peak electron density of about 30% during a solar flare (Nielsen et al. 2006). Recently, MAVEN has also detected the effect of X and M class flares in the upper thermosphere of Mars (Lee et al., 2017; Thiemann et al., 2015)

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Fig. 13.1 Dependence of total number density with altitude on Mars for non-flare and flare affected orbit on 10 September, 2017 (from Cramer et al. 2020)

The atmospheric response to X8.2 flare on the density and composition of the Mars’ thermosphere has been observed on 10 September, 2017 from MAVEN/ NGIMS (Cramer et al. 2020). The Fig. 13.1 shows the altitude dependence of total number density composed of six atmospheric gases (He, O, N2 , CO, Ar and CO2 ). The base line curve represents the mean density before/after the flare orbits. The orbits # 5715, 5716, 5717, 5718, 5719 and 5720 are the flare orbits. The total densities of six gases are plotted in Fig. 13.1 for flare and non-flare orbits on 10 September, 2017. The number densities remain unaltered due to flare effect at low altitude ~ 160 km. Above this altitude, the upper thermosphere is expanded due to flare heating and thermospheric density was increased by a factor of ~ 5. The uncertainty for the baseline total density profile is also plotted as estimated by Cramer et al. (2020).

13.2 Effects of CMEs on the Upper Ionosphere The CME affects the E-region ionosphere of Mars and increases the total electron content by a factor of 3–5 in comparison to normal conditions (Haider et al. 2009b, 2012). The physical processes of magnetic storm due to impact of CMEs

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are different on Earth and Mars. During the magnetic storm on Earth the magnetosphere is compressed by shock waves driven by CMEs and the aurora is formed due to precipitation of high energy particles into the ionosphere. Mars does not have a dipole magnetosphere. However, magnetosheath is observed in the sunlit hemisphere of Mars at about 435 km during solar quiet condition (Mitchell et al. 2000). During the magnetic storms at Mars the interplanetary shocks compress the magnetosheath to lower altitudes due to the impact of violent eruption of CMEs from the sun (Crider et al. 2005; Kallio et al. 2010; Lee et al. 2018; Haider et al. 2022). Figure 13.2 represents the proton flux distribution between 12 and 18 May 2005 measured by GOES 11 at Earth for three energies ≥ 10 meV (red color), ≥ 50 meV (black color) and ≥ 100 meV (blue color). The protons of energy ≥ 10 meV were accelerated throughout a large fraction of the heliosphere. During this event the proton flux of energy ≥ 10 meV increased by about three orders of magnitude on 15 May at 00:05:00 UT. The X-ray flare occurred on 13 May, 2005 at 17:00 h. A time series of solar X-ray flare flux distribution is shown in Fig. 13.3 (Haider et al. 2009b; Haider and Mahajan 2014). Figure 13.4 represents the measured Total Electron Content (TEC) in the E region ionosphere of Mars between 12 and 18 May 2005. The estimated TEC is also shown in this figure at flare time. In this figure 62 electron density profiles were used to obtain TEC. All the profiles were observed at nearly same latitude (65.3 °N–65.6 °N), the same SZA (82.4°–83.5°) and the same local time (14.4–4.5 LT) (Hinson et al. 1999). In this figure the effect of CME arrival at Mars after 1–2 days are marked by arrow at 14:48:28 UT and 16:46:09 UT on 16 May 2005 and at 00:32:49 UT and 04:32:10 UT on 17 May 2005 (Haider and Mahajan 2014). This process is known as magnetic storm at Mars (Haider et al. 2009b, 2012; Haider and Mahajan, 2014; Haider et al. 2022). In presence of magnetic storms the electron densities were increased by a factor of 3–5 due to CME impact in the E region ionosphere of Mars (Haider and Mahajan 2014). Fig. 13.2 Proton flux distribution between 12 and 18 May 2005 measured by GOES 11 spacecraft at Earth for three energies: ≥ 10 (red), ≥ 50 (black), and ≥ 100 meV (blue) (from Haider et al. 2009b)

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Fig. 13.3 Solar X-ray flux distributions between 12 and 18 May 2005 measured by GOES 12 spacecraft at Earth for two wavelength bands: short X-ray (XS, 0.5–4 Å, black) and long X-ray (XL, 1–8 Å, red) (from Haider et al. 2009b)

Fig. 13.4 Measured TEC in the E region of Mars’ ionosphere between 12 and 18 May 2005. Estimated result is shown by red colour (from Haider et al. 2009b)

13.3 Effects of SEPs on the Upper Ionosphere The major disturbances happen in the heliosphere due to SEP events. During this event the electron and the proton fluxes are enhanced significantly. There are mainly two types of SEP namely impulsive and gradual events, which are accelerated by solar X-ray flares and CMEs respectively (Cane et al. 1986). Impulsive events are relatively of short duration (< 1 day) while gradual events are of longer duration

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(days). There have been several efforts to describe the interaction of SEPs with the Mars’ atmosphere. Initially Leblanc et al. (2002) investigated theoretical altitude profiles of energy deposition rate in the Martian ionosphere due to SEP impact on 20 October, 1995. Later MARSIS onboard MEX observed that the electron densities were increased significantly in the upper ionosphere of Mars during SEP events (Morgan et al. 2006). ESA and NASA have developed a Mars Energetic Radiation Environment model to study the effects of SEP radiation around Mars (cf. McKennaLawlor et al. 2012). Sheel et al. (2012) investigated the response of SEP event of 29 September, 1989 at Mars. They found that SEP event enhanced the electron density by a factor of ~ 2–3 than that observed by MGS in the dayside ionosphere of Mars at altitude range 120–140 km. MAVEN has been making regular observations of the response of Mars to SEPs since November 2014. Six significant SEP events have been detected by SEP instrument onboard MAVEN during 15–23 December, 2014, 25–27 March, 5–7 May, 28 October - 9 November, 2015, 6–9 January, 2016 and 10– 23 September, 2017 (Jakosky et al. 2015; Lee et al. 2017; Ramstad et al. 2018). The diffuse auroras were observed by MAVEN in two major SEP events that occurred during 15–23 December, 2014 and 11–14 September, 2017 (Schneider et al. 2015a, b, 2018). Figure 13.5a–d shows a time series of electron spectra at energies 25, 50, 75, and 100 keV, observed by SEP instrument onboard MAVEN from 15 to 23 December, 2014. The red line represents smooth fitting obtained from ‘Smooth Data Moving Average Filter’ technique (https://mathworks.com/help/curvefit/smooth.html/). The smoothed data clearly shows the enhancement in the electron fluxes at 25, 50, 75, and 100 keV. The large enhancements in SEP electron spectra have been observed between 17 and 21 December, 2014 when diffuse auroras were observed by IUVS instrument. These spectra observed maximum electron fluxes ~ 2.4 × 104 , 1.3 × 104 , 7.0 × 103 and 5.1 × 103 cm−2 s−1 sr−1 at energies 25 keV, 50 keV, 75 keV and 100 keV respectively. The SEP peak electron fluxes are decreasing with increasing energy.

13.4 Effects of ENAs on the Upper Ionosphere Mars does not have a dipole magnetosphere. Therefore, the solar wind interacts directly with the upper ionosphere of Mars. The magnetosheath occurred at about 435 km in the sunlit hemisphere (Mitchell et al. 2000). Inside the magnetosheath, the planetary neutrals are mainly H atoms of hydrogen corona (Riedler et al. 1989). The ENAs are produced by charge exchange reactions between solar wind protons and hydrogen atoms (Galli et al. 2008; Milillo et al. 2009; Haider and Masoom 2019; Sakai et al. 2021). The combined atoms (H+ −H) can compress the magnetosheath of Mars similarly to that observed in the Earth’s magnetosphere. These H+ −H atoms have the same energy as the solar wind protons and move in the same direction as that of the fast protons just before the collisions (Haider et al. 2002, 2009a, b, 2022). In this way H+ −H atoms of high flux driven by CMEs penetrate deeply into the E

13 Mars Upper Ionospheric Disturbances

SEP electron flux (x103 cm-2 s-1 sr-1)

SEP electron flux (x103 cm-2 s-1 sr-1)

SEP electron flux (x103 cm-2 s-1 sr-1)

SEP electron flux (x103 cm-2 s-1 sr-1)

112 50 40

25 Kev smooth fitting

(a) 2014

30 20 10 0 Dec 15

Dec 17

Dec 19

Dec 21

Dec 23

50 50 KeV smooth fitting

(b) 2014

40 30 20 10 0 Dec 15

Dec 17

Dec 19

Dec 21

Dec 23

30 75 KeV smooth fitting

(c) 2014 20

10

0 Dec 15

Dec 17

Dec 19

Dec 21

Dec 23

30 100 KeV smooth fitting

(d) 2014 20

10

0 Dec 15

Dec 17

Dec 19

Dec 21

Dec 23

December 2014

Fig. 13.5 A time series of SEP electron fluxes at energies 25 (a), 50 (b)75 (c) and 100 keV (d) as observed by SEP instrument onboard MAVEN during 15–23 December, 2014. (from Masoom et al. 2019)

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region ionosphere of Mars. ASPERA-3 observed ENAs flux 1–3 × 106 cm−2 s−1 in the dayside atmosphere of Mars (Barabash et al. 2006; Futaana et al., 2006a, b, c; Galli et al., 2006a, b; Grigoriev et al. 2007). The proton aurora is observed by MEX and MAVEN during the daytime at wavelength 121.6 nm due to precipitation of ENAs (Deighan et al. 2018; Haider and Masoom 2019; Haider et al. 2022).

13.5 Ionospheric Modeling Due to Impact of X-ray Flares Several models have been developed to study the effect of solar flares in the ionosphere of Mars (Fox et al. 1996; Bougher et al. 1990; Mendillo et al. 2004; Lillis et al. 2010; Haider et al. 2012, 2016; Xu et al. 2018; Thirupathaiah et al. 2019; Shah et al. 2021). These methods are described in section 10: (10.1) continuity, momentum and energy equations, (10.2) MHD method, (10.3) Hybrid model, (10.4) two stream method, (10.5) Monte Carlo method and (10.6) AYS approach. These models have been used to study the various physical, chemical and dynamical processes operating in the Martian ionosphere. It is found that soft X-rays (0.5–90 Å) produced an E layer at altitude 100–110 km, while hard X-rays (0.5–3 Å) formed a D layer at altitude 25–30 km (Shah et al. 2021). During X-ray flare events the peak electron densities were increased by about a factor of ~ 2–5. This layer is observed in the dayside ionosphere of Mars only because X-ray radiation disappeared at night. X-ray flares are classified as A, B, C, M or X class flares according to their peak flux. These flares are measured by GOES space craft in the wavelength range 1–8 Å (Bornmann et al. (1996).

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Masoom, J., Haider, S.A., and Giuranna, M.: Response of dust on thermal emission spectra observed by Planetary Fourier3 Spectrometer (PFS) on-board Mars Express (MEX). Ind. J. Rad. Space Phys. vol 48, 38–44 (2019) McKenna-Lawlor, S., Gonçalves, P., Keating, A., et al.: Characterization of the particle radiation environment at three potential landing sites on Mars using ESA’s MEREM models. Icarus 218(1), 723–734 (2012) Mendillo, M., Pi, X., Smith, S., et al.: Ionospheric effects upon a satellite navigation system at Mars. Radio Sci. 39(2), 1–11 (2004) Mendillo, M., Withers, P., Hinson, D., et al.: Effects of solar flares on the ionosphere of Mars. Science 311(5764), 1135–1138 (2006) Milillo, A., Mura, A., Orsini, S., et al.: Statistical analysis of the observations of the MEX/ASPERA3 NPI in the shadow. Planet. Space Sci. 57(8–9), 1000–1007 (2009) Mitchell, D.L., Lin, R.P., Reme, H., et al.: Oxygen auger electrons observed in Mars’ ionosphere. Geophys. Res. Lett. 27(13), 1871–1874 (2000) Morgan, D.D., Gurnett, D.A., Kirchner, D.L., et al. (2006). Solar control of radar wave absorption by the Martian ionosphere. Geophys. Res. Lett. 33(13) Nakagawa, H., Terada, N., Jain, S. K., et al.: Vertical propagation of wave perturbations in the middle atmosphere on Mars by MAVEN/IUVS. J. Geophys. Res. Planets. 125(9), e2020JE006481 (2020) Ramstad, R., Holmström, M., Futaana, Y., et al.: The September 2017 SEP event in context with the current solar cycle: Mars express ASPERA-3/IMA and MAVEN/SEP observations. Geophys. Res. Lett. 45(15), 7306–7311 (2018) Rao, N.V., Leelavathi, V., Rao, S.V.B.: Variability of temperatures and gravity wave activity in the Martian thermosphere during low solar irradiance. Icarus, 114753 (2023) Rao, N.V., Leelavathi, V., Yaswanth, C., Rao, S.V.B.: Disentangling the dominant drivers of gravity wave variability in the Martian thermosphere. Astrophys. J. 936(2), 174 (2022) Riedler, W., Möhlmann, D., Oraevsky, V.N., et al.: Magnetic fields near Mars: first results. Nature 341(6243), 604–607 (1989) Sakai, S., Seki, K., Terada, N., et al.: Effects of the IMF direction on atmospheric escape from a mars-like planet under weak intrinsic magnetic field conditions. J. Geophys. Res. Space. Phys. 126(3), e2020JA028485 (2021) Schneider, N.M., Deighan, J.I., Stewart, A.I.F., et al.: MAVEN IUVS observations of the aftermath of the comet siding spring meteor shower on Mars. Geophys. Res. Lett. 42(12), 4755–4761 (2015a) Schneider, N.M., Deighan, J.I., Jain, S.K., et al.: Discovery of diffuse aurora on Mars. Science 350(6261), aad0313 (2015b) Schneider, N.M., Jain, S.K., Deighan, J., et al.: Global aurora on Mars during the September 2017 space weather event. Geophys. Res. Lett. 45(15), 7391–7398 (2018) Shah, S.Y., et al.: A coupled model of the D and E regions of Mars’ ionosphere for flare and non-flare electron density profiles. Icarus 361 114403 (2021) Sheel, V., Haider, S.A., Withers, P., et al.: Numerical simulation of the effects of a solar energetic particle event on the ionosphere of Mars. J. Geophys. Res. Space Phys. 117(A5) (2012) Thiemann, E.M.B., Eparvier, F.G., Andersson, L.A., et al.: Neutral density response to solar flares at Mars. Geophys. Res. Lett. 42(21), 8986–8992 (2015) Thirupathaiah, P., Shah, S.Y., Haider, S.A.: Characteristics of solar X-ray flares and their effects on the ionosphere and human exploration to Mars: MGS radio science observations. Icarus 330, 60–74 (2019) Xu, S., Fang, X., Mitchell, D.L., Ma, Y., et al.: Investigation of Martian magnetic topology response to 2017 September ICME. Geophys. Res. Lett. 45(15), 7337–7346 (2018) Zhang, S.R., Coster, A.J., Erickson, P.J., et al.: Traveling ionospheric disturbances and ionospheric perturbations associated with solar flares in September 2017. J. Geophys. Res. Space Phys. 124(7), 5894–5917 (2019)

Chapter 14

Upper Ionosphere of Mars During Low, Medium and High Solar Activity

Abstract For upper ionospheric studies, the most relevant portions of the solar spectrum are solar (EUV) and X-rays, which have sufficient energies to ionize atoms and molecules during different solar activity. The flux of photons capable of ejecting electrons from neutral ionospheric constituents is often referred to as the ionizing irradiance. To what extent these similarities and differences translate to their ionospheric responses to changing solar activity has not yet been fully determined. Studying how Mars upper ionosphere responds to changes in ionospheric parameters are important to understand during low, medium and high solar activity. In this chapter the solar activity and its effect on the upper ionosphere of Mars are discussed by using the ionospheric measuements, modeling and solar radiation. Keywords Solar activity · Ionospheric disturbances · Sunspot variability

14.1 Sunspot and Solar Activity The solar cycle has duration of 11 years on an average from solar minimum to solar maximum. The progression of solar cycle 21–25 is shown in Fig. 14.1. The sunspots are clearly visible on the sun during the solar maximum almost all the time. The solar activity is highly variable from solar minimum to solar maximum period. The sunspots are areas where the magnetic field is about 2500 times stronger than Earth’s, which is much higher than anywhere else on the sun (Nebdi 2019)Because of the strong magnetic field, the magnetic pressure increases while the surrounding atmospheric pressure decreases. This in turn lowers the temperature relative to its surroundings because the concentrated magnetic field inhibits the flow of hot plasma from the sun’s interior to the surface. At solar minimum sunspots tend to form around latitude of 30°–45° North and South of sun’s equator. As the solar cycle progresses through solar maximum, sunspots tend to appear closer to the equator around latitude of 15° (Hathaway 2015) A peak in the sunspot count is referred to as a time of solar maximum, whereas a period when a few sunspots appear is called a solar minimum. An example of recent sunspot solar cycle 24 spans the years from solar minimum in 2010 and solar © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_14

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Fig. 14.1 Progression of solar cycle 21–24 with the sunspot number (https://www.skyatnightma gazine.com/space-science/understanding-sun-science-solar-cycle/)

maximum in 2015, when about 13–100 sunspots respectively were seen (Singh and Tonk 2014). In Fig. 14.1 the maximum sunspot number is 225 during solar cycle 21. The peak sunspot numbers corresponding to solar maximum condition are decreasing gradually from 250 to 100 during solar cycle 21–25.

14.2 Effects of Solar Activity in the Upper Ionosphere During solar maximum period the electron density increased by a factor of 3–5 in the upper ionosphere of Mars (Hensley and Withers 2021). Several investigators (Breus et al. 2004; Zou et al. 2006; Fox and Yeager 2009; Withers et al. 2015; Sanchez-Cano et al. 2016, 2021; Duru et al. 2019; Tian et al. 2022) have quantified Mars’s response to changing solar irradiance at the altitude of peak electron density. Recently, Vigren and Cui (2019) have reported that the electron temperature also changes from 120 to 175 km due to solar activity. The high solar activity can also have a significant impact on the density and compositions of the Martian topside ionosphere (Mahaffy et al. 2015a, b; Thiemann et al. 2018). The MGS observations also confirmed that E peak is distinct from the F peak during the solar maximum period (Withers et al. 2021; Hensley and Withers 2021). Venkateswara Rao et al. (2014) showed that the electron density can vary with the rotation of the Sun during solar maximum condition.

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Fig. 14.2 The electron density profiles for solar minimum (MAVEN black) and solar maximum (MGS, gray) conditions (from Withers et al. 2021)

14.3 Ionospheric Measurments: Low, Medium and High Solar Activity Mariner 4 carried out ionospheric measurements during solar minimum condition, while Mariners 6 and 7 observations were made during solar maximum period. The higher solar activity explained the Mariners 6 and 7 measurements of larger electron density, greater scale heights, and higher peak altitudes. The comprehensive reviews on the Mars’ ionospheric studies covering from solar minimum to maximum period have been published by several investigators (Whitten and Colin 1974; Schunk and Nagy 1980; Mahajan and Kar 1988; Kar 1996; Nagy et al. 2004; Haider et al. 2011; Haider and Mahajan 2014; Haider et al. 2022). The radio science experiment onboard MGS has measured 5600 electron density profiles during the period 24 December 1998 to 9 June 2005, thus covering a major part of the sunspot cycle 23 (Fallows et al. 2015; Haider et al. 2016). Recently Withers et al. (2021) have compared MAVEN and MGS radio occultation profiles of radio science instruments for 23 March 2018 and 20 December 2000 during solar minimum and maximum conditions respectively. They have been found that the peak electron density is greater in the solar maximum than in the solar minimum by a factor of ~ 2 (Fig. 2).

14.4 Ionospheric Modeling: Low, Medium and High Solar Activity In the past several theoretical models of the Mars’ ionosphere have been developed and validated with the ionospheric observations during solar maximum, medium or minimum conditions (Shinagawa and Cravens 1992; Krasnopolsky 2002; Ma

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et al. 2004; Modolo et al. 2006; Brecht and Ledvina 2006; Ledvina et al. 2008; Bougher et al. 2009;. Valeille et al. 2009; Terada et al. 2009; Brain et al. 2010; Kallio et al. 2010; Matta et al. 2014; Chaufray et al. 2014; Fox 2015; Withers et al. 2015; Haider et al. 2016). Recently, Mukundan et al. (2021, 2022) have used one dimension photochemical model for the calculation of ion and electron densities in the Martian ionosphere and compared these with the NGIMS and ROES observations of MAVEN carried out during solar minimum conditions.

14.5 Solar Rotation Effects on the Martian Ionosphere The Sun rotates on its axis once in about 27 days at the equator and almost 38 days at the poles. This differential rotation causes magnetic field lines to twist resulting in the formation of active regions that release enhanced solar energy in various forms including solar EUV radiation responsible for heating in the upper ionosphere of Mars. The rotation of solar active regions produces periodicities in solar EUV flux emanating from the Sun and subsequently absorbed by Martian thermosphere (Forbes et al. 2006). The 27 day solar rotation effect in the Martian thermospheric density (CO2 , Ar, and N2 ) has been identified by analyzing five years of MAVEN observations (Hughes et al. 2022). The MAVEN has a EUV monitor and a NGIMS instrument measuring the abundance of neutrals and ions. The analyses of these data reveal that the presence of strong solar rotation effects with density increasing as EUV flux increases. These effects are strongest at higher altitudes (200–250 km) for all species, while dependence of the slope on solar zenith angle depends on the species, CO2 having a higher slope on the nightside, Ar having a higher slope on the dayside and N2 having a different behavior depending on altitude. Venkateswara Rao et al. (2014) also studied the solar rotation effects on the Martian ionosphere at high latitude (~63° N–81° N) using the electron density measured by MGS and solar XUV and EUV fluxes measured by SOHO under high, medium and low solar activity conditions. They found that the solar rotation effect is same at all altitudes though its amplitude is strongest at ionospheric F peak and has secondary enhancement at the E peak. Further the effects of solar rotations on the F peak were larger during medium solar activity than during high solar activity. Figure 14.3 represents the solar rotation effect on TEC observed by MGS during solar medium condition. This figure confirms that periodicity in TEC is due to solar rotation at ~26 days for about 70 days between between 22 March 2003 and 30 May 2003. The effect of solar rotation is also observed from MAVEN in the thermospheric density and composition of Mars. Recently Hughes et al. (2022) observed the response of CO2 , Ar, and N2 densities to 27 day solar rotation variability over 0– 7 nm, 17–22 nm, 0–45 nm and 117–125 nm spectral bands. Their results confirm the presence of strong solar rotation effects with density increasing as EUV flux increases. These effects are strongest at higher altitudes (200–250 km).

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Fig. 14.3 TEC and XUV flux variability for 22 March 2003 to May 2003 due to solar rotation (from Venkateswara Rao et al. 2014)

References Bougher, S.W., McDunn, T.M., Zoldak, K.A., Forbes, J.M.: Solar cycle variability of Mars dayside exospheric temperatures: model evaluation of underlying thermal balances. Geophys. Res. Lett. 36(5) (2009) Brain, D., Barabash, S., Boesswetter, A., et al.: A comparison of global models for the solar wind interaction with Mars. Icarus 206(1), 139–151 (2010) Brecht, S.H., Ledvina, S.A.: The solar wind interaction with the Martian ionosphere/atmosphere, Space Sci. Rev. 126, 15 (2006) Breus, T.K., Krymskii, A.M., Crider, D.H., et al.: Effect of the solar radiation in the topside atmosphere/ionosphere of Mars: Mars global surveyor observations. J. Geophys. Res. Space Phys. 109(A9) (2004) Chaufray, J.Y., Gonzalez-Galindo, F., Forget, F., et al. Three-dimensional Martian ionosphere model: II effect of transport processes due to pressure gradients. J. Geophys. Res. Planets. 119(7), 1614–163 (2014) Duru, F., Brain, B., Gurnett, D.A., et al.: Electron density profiles in the upper ionosphere of Mars from 11 years of MARSIS data: Variability due to seasons, solar cycle, and crustal magnetic fields. J Geophys. Res. Space Phys. 124(4), 3057–3066 (2019) Fallows, K., Withers, P., Gonzalez, G.: Response of the Mars ionosphere to solar flares: analysis of MGS radio occultation data. J. Geophys. Res. Space Phys. 120(11), 9805–9825 (2015) Forbes, J.M., Russell, J., Miyahara, S., et al. Troposphere-thermosphere tidal coupling as measured by the SABER instrument on TIMED during July–September 2002. J. Geophys. Res. Space Phys. 111(A10) (2006) Fox, J.L.: The chemistry of protonated species in the Martian ionosphere. Icarus, 252, 366–392 (2015) Fox, J.L., Yeager, K.E.: MGS electron density profiles: analysis of the peak magnitudes. Icarus 200(2), 468–479 (2009) Haider, S.A., Mahajan, K.K.: Lower and upper ionosphere of Mars. Space Sci. Rev. 182(1), 19–84 (2014) Haider, S.A., Batista, I.S., Abdu, M.A., et al.: Flare X-ray photochemistry of the E region ionosphere of Mars. J. Geophys. Res. Space Phys. 121(7), 6870–6888 (2016) Haider, S.A., Mahajan, K.K., Bougher, S.W., Schneider, N.M., et al.: Observations and modeling of Martian auroras. Space Sci. Rev. 218(4), 1–53 (2022)

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Haider, S.A., Mahajan, K.K., Kallio, E.: Mars ionosphere: a review of experimental results and modeling studies. Rev. Geophys. 49(4) (2011) Hathaway, D.H.: The solar cycle. Living Rev. Sol. Phys. 12(1), 1–87 (2015) Hensley, K., Withers, P.: Response of Mars’s topside ionosphere to changing solar activity and comparisons to Venus. J. Geophys. Res. Space. Phys. 126(3), e2020JA028913 (2021) Hughes, J., Gasperini, F., Forbes, J.M.: Solar rotation effects in martian thermospheric density as revealed by five years of MAVEN observations. J. Geophys. Res. Planets. 127(1), e2021JE007036 (2022) Kallio, E., Liu, K., Jarvinen, R., et al.: Oxygen ion escape at Mars in a hybrid model: high energy and low energy ions. Icarus 206(1), 152–163 (2010) Kar, J.: Recent advances in planetary ionospheres. Space Sci. Rev. 77(3), 193–266 (1996) Krasnopolsky, V.A.: Mars’ upper atmosphere and ionosphere at low, medium, and high solar activities: implications for evolution of water. J. Geophys. Res. Planets. 107(E12), 11–21 (2002) Ledvina, S.A., Ma, Y.J., Kallio, E.: Modeling and simulating flowing plasmas and related phenomena. Space Sci. Rev. 139(1), 143–189 (2008) Ma, Y., Nagy, A.F., Sokolov, I.V., Hansen, K.C.: Three-dimensional, multispecies, high spatial resolution MHD studies of the solar wind interaction with Mars. J. Geophys. Res. Space Phys. 109(A7) (2004) Mahaffy, P.R., Benna, M., Elrod, M., et al.: Structure and composition of the neutral upper atmosphere of Mars from the MAVEN NGIMS investigation. Geophys. Res. Lett. 42(21), 8951–8957 (2015a) Mahaffy, P.R., Benna, M., King, T., et al.: The neutral gas and ion mass spectrometer on the Mars atmosphere and volatile evolution mission. Space Sci. Rev. 195(1), 49–73 (2015b) Mahajan, K.K., Kar, J.: Planetary ionospheres. Space Sci. Rev. 47(3), 303–397 (1988) Matta, M., Galand, M., Moore, L., et al.: Numerical simulations of ion and electron temperatures in the ionosphere of Mars: multiple ions and diurnal variations. Icarus 227, 78–88 (2014) Modolo, R., Chanteur, G.M., Dubinin, E., Matthews, A.P.: Simulated solar wind plasma interaction with the Martian exosphere: influence of the solar EUV flux on the bow shock and the magnetic pile-up boundary. Ann. Geophysi. 24(12), 3403–3410 (2006). Copernicus GmbH Mukundan, V., Thampi, S.V., Bhardwaj, A., Fang, X.: Impact of the 2018 Mars global dust storm on the ionospheric peak: a study using a photochemical model. J. Geophys. Res. Planets. 126(4), e2021JE006823 (2021) Mukundan, V., Thampi, S.V., Bhardwaj, A.: M3 electron density layer in the dayside ionosphere of Mars: Analysis of MAVEN ROSE observations. Icarus 384, 115062 (2022) Nagy, A., Ma, Y., Sokolov, I.: 3D, multi-species, high spatial resolution MHD studies of the solar wind interaction with Mars. 35th COSPAR Scientific Assembly. vol. 35. (2004) Nebdi H (2019) Space weather and link to climate change. In: Handbook of Research on Global Environmental Changes and Human Health, pp. 1–20. IGI Global Sánchez-Cano, B., Lester, M., Witasse, O., et al.: Solar cycle variations in the ionosphere of Mars as seen by multiple Mars Express data sets. J. Geophys. Res. Space Phys. 121(3), 2547–2568 (2016) Sánchez-Cano, B., Lester, M., Cartacci, M., et al. Ionosphere of Mars during the consecutive solar minima 23/24 and 24/25 as seen by MARSIS-Mars express. Icarus, 114616 (2021) Schunk, R.W., Nagy, A.F.: Ionospheres of the terrestrial planets. Rev. Geophys. 18(4), 813–852 (1980) Shinagawa, H., Cravens, T.E.: The ionospheric effects of a weak intrinsic magnetic field at Mars. J. Geophys. Res. Planets. 97(E1), 1027–1035 (1992) Singh, A.K., Tonk, A.: Solar activity during first six years of solar cycle 24 and 23: a comparative study. Astrophys. Space Sci. 353(2), 367–371 (2014) Terada, N., Kulikov, Y.N., Lammer, H., et al.: Atmosphere and water loss from early Mars under extreme solar wind and extreme ultraviolet conditions. Astrobiology 9(1), 55–70 (2009) Thiemann, E.M.B., Andersson, L., Lillis, R., et al.: The Mars topside ionosphere response to the X8 2 solar flare of 10 September 2017. Geophys. Res. Lett. 45(16), 8005–8013 (2018)

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Chapter 15

Characteristics of Martian Ionopauses

Abstract The Martian ionopause is not always formed and when it is formed, it has no static boundary (Schunk and Nagy in Ionospheres: physics, plasma physics, and chemistry. Cambridge University Press, 2009; Duru et al. in J Geophys Res Space Phys 114(A12), 2009; Chu et al. in Geophys Res Lett 46(17–18):10257– 10266, 2019). The ionopause is narrow and its altitude is high, when the solar wind dynamic pressure is low relative to the ionospheric thermal pressure (Withers et al. in Planet Space Sci 20:24–34, 2016). When the solar wind dynamic pressure is relatively high, the ionospheric thermal pressure alone is not sufficient to balance it and the ionosphere becomes magnetized with large scale horizontal fields. These observations have lacked of simultaneous measurements of solar wind, magnetic field and the ionosphere. MAVEN carried a suite of plasma and magnetic field instruments viz. MAG (Connerney et al. in Geophys Res Lett 42(21):8819–8827, 2015), NGIMS (Mahaffy et al. in Space Sci Rev 195(1):49–73, 2015), SWIA (Halekas et al. in Geophys Res Lett 42(21):8901–8909, 2015), and LPW (Andersson et al. in Space Sci Rev 195(1):173–198, 2015). The data of these instruments have been used to study the full characterization of the Martian ionopauses. Keywords Ionopause · Solar wind · Magnetic field

15.1 Earlier Measurements of Ionopauses The ionosphere of Mars was first observed by RO experiment onboard Mariner 4 (Kliore et al. 1965). The Mariner 4 was followed by Mariner 6, 7, and 9, Mars 2, 3 and Viking 1, 2 (Kolosov et al. 1972, 1973, 1976; Kliore et al. 1973; Vasilev et al. 1975; Savich and Samovol 1976; Fjeldbo et al. 1977). After about two decades later, MGS observed 5600 electron density profiles between 24 December 1998 and 9 June 2005 at SZA 71°–89° (Withers et al. 2008). Later Radio Science (RS) experiment onboard MEX observed about 900 electron density profiles from 2004 to 2018 at SZA 50°–100° during solar cycle 23 and 24 (cf. Withers et al. 2016; Patzold et al. 2016; Dubinin et al. 2017). Most of these electron density profiles were observed

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between ~70 km and ~350 km. However, only few profiles were extended from 60 km to about 800 km altitude in the southern hemisphere. Nearly 90% of electron density profiles from these measurements showed no abrupt drop at the ionopause altitude. These electron densities were decreasing with altitude and gradually approached the measurement noise level at the top of the Mars’ ionosphere (Ness et al. 2000; Vogt et al. 2016). However, ionopause-like structures were observed in ~10% profiles of Mariner 9 and Viking 1/2 (Fjeldbo et al. 1977; Kliore 1992). These electron density profiles represented the ionopause altitudes at about 300–350 km, where the electron density is suddenly reduced to ~102 cm−3 . In addition to RS experiment MEX also carried out MARSIS instrument, which provided electron density profiles well above the main ionospheric peak (Gurnett et al. 2005). MARSIS observed a steep transient electron density gradient in the topside ionosphere of Mars (Duru et al. 2009). Such sharp gradients were not detected in any of the earlier measurements of Mariners, Vikings, MGS and MEX because radio occultation experiments were unable to resolve the electron densities less than 103 cm−3 . The electron density values as low as ~20 cm−3 have been observed from MARSIS instrument. Later Chu et al. (2019) also used MARSIS data for the study of the ionopause of Mars. They found that the ionopause is located on an average at an altitude of 363 ± 65 km and it has a weak dependence on SZA but varies with solar EUV fluxes on an annual and solar time scales. These ionopause structures were observed at about 400 km in presence of magnetic field of ~30–40 nT. In the mini-magnetosphere region of Mars the solar wind interaction with the ionosphere is not direct allowed and this prevents the ionopause formation (Chu et al. 2019).

15.2 Ionopause Measurements from MAVEN Vogt et al. (2015) and Sánchez-Cano et al. (2020) reported preliminary results of ion density and electron density measurements from NGIMS and LPW instruments respectively. Vogt et al. (2015) selected ionopause, where ion density decreased suddenly by at least a factor of 10 over an altitude range of at most 30 km. They have found ionopause-like gradients at altitude range from 270 to 464 km in 45 of the 84 orbits of MAVEN during the months of October 2014, April 2015 and May 2015. These altitudes are lower than those reported by Duru et al. (2009) (i.e. 450–500 km). Duru et al. (2009) reported sharp density gradients in only about 10% of MARSIS orbits. It should be noted that the available evidence of the Martian ionopause identified by Vogt et al. (2015) do not confirm that the crustal magnetic field strength or solar wind dynamic pressure influences on the ionopause altitude, although the previous studies have reported that these factors appear to influence the presence or absence of an ionopause (Withers et al. 2016; Chu et al. 2019). Thus, the ionopause of Mars is not well understood.

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Sánchez-Cano et al. (2020) studied 114 deep dip orbits of MAVEN between periapsis and 600–1000 km. They have identified steep ionopause-like boundary, where electron density decreases and temperature increases suddenly with altitude below 400 km. They have also evaluated the roles of ionospheric and magnetic field pressures on the ionopase formation as well as the presence of solar wind particles H+ , down to the location of the ionopause. They argued that the ionopause is formed where the total ionospheric pressure (magnetic + thermal) is equal to the upstream solar wind dynamic pressure. These ionopause are rarely formed over a strong crustal magnetic field because the magnetic pressure is larger and stands off to the solar wind dynamic pressure more often. They have found that 45% of MAVEN electron density profiles exhibited an ionopause.

15.3 Magnetic Pile-Up Boundary The Magnetic Pile-up Boundary (MPB) is identified where the magnetic field shootsup quickly. In this region two broad peaks were observed in the magnetic fields (Sometimes these peaks are not clearly separated). The Cavity Boundary (CB) is located between these two magnetic field peaks. The total magnetic field is equal to the square root of the sum of BH 2 and Br 2 , where BH and Br are horizontal and radial magnetic fields respectively. In Fig. 15.1d the total magnetic field is dominated by the horizontal magnetic field because the radial magnetic field is very small (see Figs. 15.1b–c). The horizontal magnetic fields are suddenly increased by a factor of 3–4 within the magnetic pile-up boundary. This confirms that the IMF draping in the magnetic-pile-up boundary is mainly horizontal. In this figure two broad peaks and a drop in the time series of the horizontal magnetic fields have been observed inside the magnetic pile-up boundary. Figure 15.1a represents the spacecraft positions with LT and SZA, while Fig. 15.1e, f show the variations of H+ density and electron density, Ne respectively observed by MAVEN on 13 October, 2014.

15.4 Low, Mid and High Altitude Ionopauses Recently it has been observed that the electron densities in the daytime ionosphere of Mars are suddenly reduced at low, mid, and high altitudes range ~ 350–450 km, 500–700 km, and 800–1000 km respectively (Haider et al. 2023) in presence of horizontal magnetic field of high strength during the inbound or outbound orbit of MAVEN within the magnetic pile-up boundary. These ionopause boundaries are controlled by the horizontal magnetic field within the magnetic pile-up region at SZA ~60°–106°.The horizontal magnetic field inhibits the upward diffusion of plasma and reduces the electron densities within these altitudes range by 1–2 orders of magnitude

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Fig. 15.1 A time series of the spacecraft orbit (a), horizontal magnetic field (b), radial magnetic field (c), total magnetic field (d), H+ density, (e), and electron density (f) respectively, when the MAVEN was passing from Magnetic Pile-Up Boundary (MPB) and Cavity Boundary (CB) on 13 October, 2014. The horizontal magnetic field is of high strength, while radial magnetic field is low inside the MPB region. The electron density represents a broad peak in the CB region (from Haider et al. 2023)

15.4 Low, Mid and High Altitude Ionopauses

02:41-02:52 UT (Inbound) 02:52-03:05 UT (Outbound)

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Fig. 15.2 The altitude profiles of electron densities during the daytime and nighttime ionosphere on 1 May 2018. The deep ionopause-like structure is seen in the electron density of the daytime ionosphere (low altitude ionopause: black colour). The ionopause-like structure is not observed in the electron densities of the nighttime ionosphere (red colour) (from Haider et al. 2023)

as shown in Figs. 15.2, 15.3, and 15.4. In Fig. 15.2 the electron density profiles during inbound (02:41–02:52 UT) and outbound (02:52–03:05 UT) orbits # 6975 on 1 May, 2018 are shown. In Fig. 15.3 the electron density profiles during inbound (05:50–06:08 UT) and outbound (06:08–06:22 UT) orbits # 138 on 24 October, 2014 are shown. In Fig. 15.4 the electron density profiles during inbound (20:30–20:55 UT) and outbound (20:55–21:14 UT) orbits # 134 on 23 October, 2014 are shown. In presence of a strong radial/local field the vertical diffusion of plasma is allowed above 200 km (Ma et al. 2004). Haider et al. (2023) suggested that in absence of radial magnetic field inside the magnetic pile-up boundary, the electrons are transported from dayside to nightside ionosphere across the terminator by a horizontal plasma flow velocity. Therefore, the ion/electron densities are nearly same in the nighttime and daytime ionosphere below the ionopause altitude within the magnetic pile-up boundary.

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Fig. 15.3 The altitude profiles of electron densities during the daytime and nighttime ionosphere on 24 October 2014. The deep ionopause-like structure is seen in the electron density of the daytime ionosphere (Mid-altitude ionopause: red colour). The ionopause-like structure is not observed in the electron densities of the nighttime ionosphere (black colour) (from Haider et al. 2023) 1200 20:30-20:55 UT (Inbound) 20:55-21:14 UT (Outbound)

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Fig. 15.4 The altitude profiles of electron densities during the daytime and nighttime ionosphere on 23 October 2014. The deep ionopause-like structure is seen in the electron density of the daytime ionosphere (high altitude ionopause: red colour). The ionopause-like structure is not observed in the electron density ptofile of the nighttime ionosphere (black colour). (from Haider et al. 2023)

15.5 Basic Equations for Ionopause Modeling

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15.5 Basic Equations for Ionopause Modeling Following Shinagawa and Cravens (1989), a set of equations for the ionopause modeling controlled by the induced magnetic field in terms of continuity equations, momentum equations and divergence term of horizontal velocity are given below: Continuity Equations ∂qi ∂ Wi ∂ = pi − qi li − (qi Vi ) − qi ∂t ∂z ∂y

(15.1)

Momentum equations     ∂qi 1 ∂ Tp ∂ αi ∂ Ti Te ∂qe 1 ∂ (15.2) (qi V i ) = Di + qi + + + ∂z ∂z ∂z Ti ∂z Ti ∂ z Ti qe ∂ z Hi Divergence term of horizontal Velocity  1/2  νin ∂ Wi 4B2H /μ0 = −1 + 2 2 ∂y 2 m i qe νin rm

(15.3)

where subscript i denotes ith ion species, qi is ion density, V i is the vertical velocity, Hi is the scale height, Tp is total temperature equal to Ti + Te , Ti and Te are ion and electron temperature respectively, Di is ion diffusion coefficient, mi is the mass of ion, αi is the ion thermal diffusion coefficient, pi and li are the production and loss coefficients respectively, Wi is the horizontal velocity in Y direction, rm is the distance between the centre of Mars and ionosphere, BH is intensity of the measured magnetic field in horizontal direction, νin is the ionization frequencies and μo is the permeability of vacuum. In order to solve the above equation, lower and upper boundary conditions are required: (1) photochemical equilibrium condition is imposed on the ion densities at the lower boundary (150 km), where the vertical velocity Vi is kept zero, and (2) at the upper boundary the vertical velocity is again set to be zero at the ionopause altitude, which is observed at 500 km on 10 and 14 September, 2017. It should be noted that ionopause altitudes are changing with solar wind dynamic pressure. Since Eq. (15.1) do not include a spatial derivative of the horizontal ion flux, it is not necessary to specify the boundary condition for Wi . In deriving horizontal divergence term in Eq. (15.3), Shinagawa and Cravens (1989) used Newtonian distribution, where magnetic field pressure varies as cos2 χ in the horizontal direction (where χ is a SZA, which is zero at the sub-solar point). This equation represents the approximate ion loss term due to the divergence of the horizontal velocity in the continuity Eq. (15.1). This loss term reduced the electron density by ~70% near the ionopause in presence of horizontal magnetic field of ~30–40 nT. It should be noted that the effect of vertical velocity on the horizontal velocity i.e. Vi (òWi /òz) is negligible in Eq. (15.1). Further, Wi (òqi /òy) is also neglected in this equation

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because Wi is assumed to be very small below the ionopause. The diffusion coefficient Di is calculated by ion-neutral polarization interactions with each neutral, ion-neutral resonant charge exchange and ion-ion Coulomb interaction. The thermal diffusion coefficient αi can be calculated using the formula given by St Maurice and Schunk (1977). The chemical reactions are important below 200 km. Above this altitude, the plasma transport is needs to be included in the photochemical equilibrium model. This is because of the fact that chemical life time is very large at altitudes 200 km and the above partial differential equations can be solved by using the finite difference method. At high SZA (near the terminator), the horizontal motion on the ion densities might be significant near the ionopause. This effect is incorporated including Eq. (15.3) in Eq. (15.1). The computer code of Shinagawa and Cravens (1989) has been used by Haider and Oyama (2002) for the analysis of radio occultation electron density profiles obtained from MGS and Viking spacecrafts. Haider et al. (2023) have used Shinagawa and Cravens (1989) code again for the calculation of low altitude ionopauses (~400– 450 km) on 10 and 14 September, 2017. In Fig. 15.5a 10 September 2017, inbound (17:30–17:42 UT) and outbound (17:42–17:52 UT) and Fig. 15.5b 17 September 2017, inbound (19:02–19:14 UT) and outbound (19:14–19:26 UT) electron density observations are shown. The estimated results are compared with the electron density measurements carried out by LPW instrument (see Fig. 15.5a, b). There is a good agreement between these observations and the estimated results. The modeling of mid- and high altitude ionopauses were not carried out because the neutral densities were not measured by NGIMS at altitude beyond 500 km (Haider et al. 2023). The mid-and high altitude ionopauses occurred between altitude range ~ 500–700 km and ~800–1000 km respectively. This code is run for three ion species, CO2 + , O2 + and O+ . The neutral densities of these gases are taken from NGIMS for 10 and 14 September, 2017. The photoionization rates are calculated at SZA ~70° by using the method of Haider and Oyama (2002) using solar EUV radiations between 1 and 102.57 nm. The major ion O+ is produced by photoionization process above 200 km because the atomic oxygen is a dominant species in the upper ionosphere of Mars. The chemical scheme of Shinagawa and Cravens (1989) is used in the above model. The neutral, ion and electron temperatures necessary to calculate the electron density are taken from Haider et al. (2009) for averaged condition. The νin is calculated from the photoionization rate of O+ after dividing it with the neutral density of atomic oxygen (Virgen and Galand 2013).

References

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Fig. 15.5 The electron density profiles observed on 10 September (a) and 14 September 2017 (b) are compared with the model results. The inbound profiles represent the ionopause-like structures. The ionopause-like structures are not observed in the outbound electron density profiles during the nighttime (red colour) (from Haider et al. 2023)

References Andersson, L., Ergun, R.E., Delory, G.T., et al.: The Langmuir probe and waves (LPW) instrument for MAVEN. Space Sci. Rev. 195(1), 173–198 (2015) Chu, F., Girazian, Z., Gurnett, D.A., Morgan, D.D., Halekas, J., et al.: The effects of crustal magnetic fields and solar EUV flux on ionopause formation at Mars. Geophys. Res. Lett. 46(17–18), 10257–10266 (2019) Connerney, J.E., Espley, J.R., DiBraccio, G.A., et al.: First results of the MAVEN magnetic field investigation. Geophys. Res. Lett. 42(21), 8819–8827 (2015) Dubinin, E., Fraenz, M., Pätzold, M., et al.: Martian ionosphere observed by Mars Express 2 Influence of solar irradiance on upper ionosphere and escape fluxes. Planet Space Sci. 145, 1–8 (2017) Duru, F., Gurnett, D.A., Frahm, R.A., Winningham, J.D., Morgan, D.D., Howes, G.G.: Steep, transient density gradients in the Martian ionosphere similar to the ionopause at Venus. J. Geophys. Res. Space Phys. 114(A12) (2009) Fjeldbo, G., Sweetnam, D., Brenkle, J., et al.: Viking radio occultation measurements of the Martian atmosphere and topography: primary mission coverage. J. Geophys. Res. 82(28), 4317–4324 (1977) Gurnett, D.A., Kirchner, D.L., Huff, R.L., et al.: Radar soundings of the ionosphere of Mars. Science 310(5756), 1929–1933 (2005) Haider, S.A., Durga K.P., Shah, S.Y.: The magnetically controlled ionopause boundary observed by LPW onboard MAVEN within magnetic pile-up region of Mars. Icarus 394, 115423 (2023) Haider, S.A., Oyama, K.I.: Calculated electron flux and densities at 10–1000 eV in the dayside Martian ionosphere: comparison with MGS and Viking results (2002) Haider, S.A., et al.: D, E, and F layers in the daytime at high-latitude terminator ionosphere of Mars: comparison with Earth’s ionosphere using COSMIC data. J. Geophys. Res. Space Phys. 114(A3) (2009) Halekas, J.S., Lillis, R.J., Mitchell, D.L.: MAVEN observations of solar wind hydrogen deposition in the atmosphere of Mars. Geophys. Res. Lett. 16;42(21), 8901–8909 Kliore, A.J.: Radio occultation observations of the ionospheres of Mars and Venus. Washington DC Am. Geophys. Union Geophys. Monogr.ser. 66, 265–276 (1992) Kliore, A., Cain, D.L., Levy, G.S.D., et al.: Occultation experiment: results of the first direct measurement of Mars’s atmosphere and ionosphere. Science 149(3689), 1243–1248 (1965)

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Kliore, A.J., Fjeldbo, G., Seidel, B.L., Sykes, M.J., Woiceshyn, P.M.: S band radio occultation measurements of the atmosphere and topography of Mars with Mariner 9: extended mission coverage of polar and intermediate latitudes. J. Geophys. Res. 78(20), 4331–4351 (1973) Kolosov, M.A., Yakovlev, O.I., Kruglov, Y.M., Trusov, B.P., et al.: Atmosphere of Mars from Mars 2 radio occultation measurement. Radiotekhn. Elektron. 17(12), 2483–2490 (1972) Kolosov, M.A., Savich, N.A., Azarkh, S.L., Aleksand, Y.N., et al.: Results of 2-frequency investigation of mars ionosphere with spacecraft mars-2. Radiotekhnika I Elektronika. 18(10), 2009–2014 (1973) Kolosov, M.A., Yakovlev, O.I., Yakovleva, G.D., Efimov, A.I., et al.: Results of investigations of Martian atmosphere by the method of radio occultation by means of space probes Mars-2, Mars-4, and Mars-6. Kosm. Issled 13(1), 54–59 (1976) Ma, Y., Nagy, A.F., Sokolov, I.V., Hansen, K.C.: Three-dimensional, multispecies, high spatial resolution MHD studies of the solar wind interaction with Mars. J. Geophys. Res. Space Phys. 109(A7) (2004) Mahaffy, P.R., Benna, M., King, T., et al.: The neutral gas and ion mass spectrometer on the Mars atmosphere and volatile evolution mission. Space Sci. Rev. 195(1), 49–73 (2015) Maurice, J.P., Schunk, R.W.: Diffusion and heat flow equations for the mid-latitude topside ionosphere. Planet Space Sci. 25(10), 907–920 (1977) Ness, N.F., Acuña, M.H., Connerney, J.E.P., et al.: Effects of magnetic anomalies discovered at Mars on the structure of the Martian ionosphere and solar wind interaction as follows from radio occultation experiments. J. Geophys. Res. Space Phys. 105(A7), 15991–16004 (2000) Pätzold, M., et al.: Mars express 10 years at Mars: Observations by the Mars express radio science experiment (MaRS). Planetary. Space Sci. 127 44–90 (2016) Sánchez-Cano, B., Narvaez, C., Lester, M., Mendillo, M., et al.: Mars’ ionopause: a matter of pressures. J. Geophys. Res. Space Phys. 125(9), e2020JA028145 (2020) Savich, N.A., Samovol, V.A.: The night-time ionosphere of Mars from Mars 4 and Mars 5 dualfrequency radio occultation measurements. Space Res. XVI, 1009–1011 (1976) Schunk, R., Nagy, A.: IonOspheres: Physics, Plasma Physics, and Chemistry. Cambridge University Press (2009) Shinagawa, H., Cravens, T.E.: A one-dimensional multispecies magnetohydrodynamic model of the dayside ionosphere of Mars. J. Geophys. Res. Space. Phys. 94(A6), 6506–6516 (1989) Vasilev, M., Vyshlov, A.S., Kolosov, M.A., Savich, N.A., et al.: Preliminary results of dualfrequency radio-occultation measurements of the Martian ionosphere by Mars spacecraft in 1974. Kosmicheskie Issledovaniia. 13, 48–53 (1975). Vigren, E., Galand, M.: Predictions of ion production rates and ion number densities within the diamagnetic cavity of Comet 67P/Churyumov–Gerasimenko at perihelion. Astrophys J 772(1), 33 (2013) Vogt, M.F., Withers, P., Mahaffy, P.R., Benna, M., et al.: Ionopause-like density gradients in the Martian ionosphere: a first look with MAVEN. Geophys. Res. Lett. 42(21), 8885–8893 (2015) Vogt, M.F., Withers, P., Fallows, K., et al.: Electron densities in the ionosphere of Mars: a comparison of MARSIS and radio occultation measurements. J. Geophys. Res. Space Phys. 121(10), 10–241 (2016) Withers, P., Matta, M., Lester, M., Andrews, D., Edberg, N.J., et al.: The morphology of the topside ionosphere of Mars under different solar wind conditions: results of a multi-instrument observing campaign by Mars Express in 2010. Planet Space Sci. 20, 24–34 (2016) Withers, P., Mendillo, M., Hinson, D.P., Cahoy, K.: Physical characteristics and occurrence rates of meteoric plasma layers detected in the Martian ionosphere by the Mars Global Surveyor Radio Science Experiment. J. Geophys. Res. Space Phys. 113(A12) (2008)

Chapter 16

Aurora and Airglow on Mars

Abstract Aurora is observed throughout the solar system (Clarke et al. 2004). It occur where the charged particles have access to an atmosphere along magnetic field lines of Earth and giant planets (Jupiter, Saturn, Uranus and Neptune) (Mcllwain 1960). Auroras on giant planets have been observed from ground based observations, Earth orbiting satellites (e.g. International Ultraviolet Explorer (IUE), Hubble Space Telescope (HST) and Roentgensatellite (ROSAT)), flyby spacecraft (e.g. Voyager 1 and 2) and orbiting spacecraft plateforms (e.g. Galileo) at X-ray, UV, Visible, IR and radio wavelengths. UV, Visible and IR auroras are atmospheric emissions produced or initiated when ambient atmospheric species are excited through the collision with precipitating particles, while radio and X-ray auroras are beam emissions, produced by precipitating species themselves. Aurora is also observed on Venus and Mars. Mars has localized crustal magnetic field. Venus has no crustal field. The Earth’s type auroras are not observed on Venus and Mars. The auroras on Venus and Mars can not be seen from nacked eyes. Recently three kinds of aurora (1) discrete aurora, (2) proton aurora, and (3) diffuse aurora, have been detected on Mars from the Imaging Ultraviolet Spectrograph (IUVS) instrument (Schneider et al. in Science 350(6261):aad0313, 2015; Haider et al. in Planet Space Sci 212:105424, 2022, Icarus (in press), Space. Sci. Rev. 218(4):1–53. The discrete aurora (Bertaux et al. in Nature 435:790–794, 2005) has been observed near the crustal magnetic field lines in southern hemisphere of Mars. This aurora is produced in the nighttime atmosphere due to precipitation of energetic electrons. The proton aurora (Deighan et al. in Nat Astron 2:802–807, 2018) is observed in the daytime due to energetic proton precipitation into the Martian atmosphere. During this event, Lyman-α limb profiles are enhanced at altitude between 120 and 150 km. The diffuse aurora (Schneider et al. in Science 350(6261):aad0313, 2015) is observed in the nighttime due to precipitation of SEP electrons down up to 1 microbar altitude. This aurora is neither restricted to location nor linked to the magnetic field. It is globally distributed and is closely correlated to solar wind velocity Keywords Aurora · Airglow · Excitation processes

© Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_16

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Figure 16.1 shows the comparison of field lines configuration for diffuse and discrete auroras on Earth and Mars. This figure represents very different magnetic field lines structure of Mars and Earth. The discrete aurora on Earth occurs at high latitude region of closed field lines connected to the planet on both ends. A similar region also occurs on Mars in the mini-magnetosphere forming the discrete aurora. Both auroras are powered by the electrons accelerated from the sun. These electrons collide with the Earth’s atmosphere along the open field lines in the auroral oval. The magnetic field lines of Mars are open or draped which allow ENAs and SEP electrons to precipitate into the atmosphere during solar storms. These particles are 100 times more energetic than the magnetospheric electrons (Schneider et al. 2015). The IMF field lines are mostly open and cover much of the planet. Therefore diffuse aurora on Mars occurs everywhere on the planet (Schneider et al. 2015; Haider et al. 2022). The observations of discrete, proton and diffuse auroras are described below: The day and night airglow are produced due to different processes. The dayglow arises when the solar radiation interacts with the daytime atmosphere. The atoms and molecules in the Martian atmosphere absorb some of the sunlight, which excites them temporarily until they release the additional energy as light either at the same or lower frequency than the absorbed light. This emission is much weaker than the light scattered from the Sun so we cannot see dayglow with the nacked eye. The nightglow on Mars occurs due to recombination of molecules that have been broken apart by solar radiation during the day. The light is emitted when excited atoms or molecules return to their original unexcited state during the night.

16.1 Discrete Aurora The discrete aurora is produced in the mini-magnetosphere of Mars due to precipitation of electrons in the nighttime along the magnetic field lines, connecting to the solar wind. This aurora was observed at an apparent tangent altitude of 19 km, when the local time was nearly 21 h and the longitude and latitude were 198.4° and 46.3° respectively (Bertaux et al. 2005). Figure 16.2a represents nighttime limb observations of H Lyman α (121.6 nm) and NO bands (190–270 nm) carried out by SPICAM UV spectrometer onboard MEX. In Fig. 16.2b auroral peaks of NO bands (181–298 nm) are shown for five bins between 533 and 540 s. This peak is confined at altitude range 60–80 km and is more intense at large southern latitudes. These measurements were carried out in the mini-magnetosphere of Mars, where field lines are nearly open and are vertical (Lundin et al. 2006; Mitchell et al. 2007). Discrete aurora was also observed by Imaging Ultraviolet Spectrograph (IUVS) onboard MAVEN. Figure 16.3 shows topside and bottomside images of the discrete aurora in the orbit number 5738 for the dayside and nightside during space weather event, which occurred between 13 and 15 September 2017 (Scneider et al. 2018). The letters indicate the locations of confirmed auroral emissions. Near the strongest crustal fields the occurrence rate of discrete aurora is ~25% of the observations. These auroras are also distributed globally in about 1% of the observations in the

16.1 Discrete Aurora

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Fig. 16.1 The magnetic field geometry for diffuse and discrete auroras on Earth (top curve) and Mars (bottom curve) Mars lacks of global dipole magnetic field due to cooling of its core (from Schneider et al. 2015; Haider et al. 2022)

regions of weak or no crustal fields. Discrete aurora is forty times stronger during the highest magnetic field regions than at the rest of the planet (Haider et al. 2022). During these events the strong electron precipitation is measured by MAVEN (Soret et al. 2021).

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Fig. 16.2 a Limb spectra observed by SPICAM between 450 to 750 s. The altitudes of the Mars Nearest Point (MNP) are shown at the rightside. This figure shows H Lyman α emission at 121.6 nm and well structured band (190–270 nm) of NO. The intensity in ADU (Analogue Digital Units) per pixel is colour-coded. b There is a sharp auroral peak separated from NO spectrum. The averaged signal intensity between 181 and 298 nm is given in ADU with respect to 200–900 s (from Bertaux et al. 2005)

16.1 Discrete Aurora

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Fig. 16.3 Top: Apoase image of Mars orbit 5738 shows the daylight side middle UV colors maped up to visible colors. Bottom: Apoase image of Mars orbit 5738 shows the nightside maped. The letters indicate the locations of confirmed emissions and the instrumental artifact (from Schneider et al. 2018; Haider et al. 2022)

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16.2 Proton Aurora The hydrogen corona is observed upto several thousand Kilometer on Mars (Chaufray et al. 2007). The solar wind protons interact with this corona through charge exchange and thereby produce a beam of ENA at solar wind speeds (Kallio and Janhunen 2001; Kallio and Barabash 2001; Gunell et al. 2006; Halekas et al. 2015). Figure 16.4 represents the mechanism of Martian proton aurora originating from solar wind charge exchange (Deighan et al. 2018; Hughes et al. 2019; Haider et al. 2022). In this process fast hydrogen atoms are produced due to charge exchange between solar wind protons (H+ ) and hydrogen corona. The fast hydrogen atoms penetrate deep in the Martian atmosphere and deposit their energy in the lower thermosphere at altitudes ~110–130 km resulting in ionization, dissociation, heating, excitation and auroral emissions. The proton aurora is produced by Lyman α (121.6 nm) due to emission of fast (excited) hydrogen atoms from the upper (2 p) state to the ground state (1 s) (Gérard, et al. 2019).

Fig. 16.4 The mechanisms of proton aurora on Mars due to solar wind charge exchange. The solar wind protons (H+ ) are deflected by the Martian bow shock (yellow lines). The neutral hydrogen in the corona (blue) becomes energetic neutral atoms (shown as proton surrounded by a shell (cyan) containing an electron (white dot). These neutral particles freely pass through the bow shock while retaining the original kinetic energy and direction of the solar wind (green arrow) and deposit their energy in the Martian thermosphere. A fraction of this energy is emitted from the energetic hydrogen as Ly α proton (from Deighan et al. 2018; Haider et al. 2022)

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The proton aurora on Mars was first observed in the daytime by Ritter et al. (2018) from SPICAM UV instrument onboard MEX. Later Deghan et al. (2018) also observed proton aurora during the daytime from IUVS instrument onboard MAVEN. In Fig. 16.5 proton aurora observed by IUVS is demonstrated for orbits 985 and 986 on 3 April 2015 at 07:33 LT and 12:04 LT respectively. In this observation the UV dayglow emissions due to CO2 ionization and its dissociation into CO and O bands (135.6 nm and 130.4 nm) are also seen. In the orbit # 985 the Lyman α represents nearly a flat profile due to multiple scattering. After about 4.5 h the Lyman α profile developed a strong peak of proton aurora in orbit # 986. In Fig. 16.6 the auroral emissions for proton auroras are plotted at different SZA. The brightest peak of this aurora occurred at SZA ≤ 40°. These auroras disappeared near the terminator. The emission intensities of this aurora represent a broad peak at ~ 130 km.

Fig. 16.5 Altitude profiles of Lyman α intensity during orbits # 985 and 986 on 3 April 2015 (from Deighan et al. 2018; Haider et al. 2022)

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Fig. 16.6 Lymal α brightness with SZA variation during orbit # 985 and 986 on 3 April 2015 (from Deighan et al. 2018; Haider et al. 2022)

16.3 Diffuse Aurora Schneider et al. (2015) observed diffuse aurora of CO2 + (B2 ∑u+ − X2 πg ) Ultraviolet Doublet (UVD) band at 289 nm from IUVS instrument in the nighttime atmosphere of Mars due to precipitation of SEP electrons at 1 microbar altitude. This aurora is neither connected to the magnetic field nor restricted to any fixed location (in terms of latitude and coordinate). Figure 16.7 represents relative spectra observed by IUVS instrument at Mars (Schneider et al. 2015, Haider et al. 2022). The top spectrum is created by solar EUV radiation. The bottom spectrum shows a nightglow feature of NO. The middle spectrum shows a nighttime auroral feature of CO2 + UVD, which are more pronounced than nightglow spectrum of NO. The auroral emissions are ~ 20 times weaker than the dayglow emissions. The diffuse aurora on Mars occurred on a global scale spanning over all longitudes in the northern hemisphere for the period from 17 to 21 December 2014 during orbits # 418 to 444. These observations were carried out in the nighttime at latitudes from 35 to 70o N with local time varying from 00:30 to 05:00.

16.3 Diffuse Aurora

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Fig. 16.7 The IUVS observations in the upper atmosphere of Mars. The top spectrum shows tree dayglow emissions from ionization and dissociation products of CO2 created by solar EUV radiation. The bottom spectrum shows nightglow NO emission. The middle spectrum shows auroral emissions at nighttime. (from Schneider et al. 2018; Haider et al. 2022)

Recently Schneider et al. (2018) have also observed 25 times brighter diffuse aurora in the nighttime from IUVS instrument onboard MAVEN during space weather events that occurred in September 2017. Figure 16.8 shows nightside image taken at the start and at the peak of the space weather events. In the left image Mars’ disk shows uniform background consistent with the instrument noise and the limb shows a faint but significant brightening. The right side image taken at the peak of these space weather events shows a pronounced auroral emission during entire nightside of Mars. The left side image was taken on 12 September 2017 at 07:24 UTC during orbit # 5726. The right side image was taken on 13 September 2017 at 05:34 UTC during orbit # 5731. Both diffuse auroral emissions observed by Schneider et al. (2015, 2018) are consistent with high energy particle precipitation.

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Fig. 16.8 Mid-ultraviolet images of Mars’ aurora during nighttime: Left image at the start of space weather event on 12 September 2017 at 07:24 UTC and right image at the peak of space weather event on 13 September 2017 at 05:34 UTC (from Schneider et al. 2018; Haider et al. 2022)

16.4 Dayglow on Mars There have been a number of observations on the dayglow from Mariners 6, 7, and 9 spacecraft (Barth et al., 1971; Stewart 1972; Barth et al. 1972; Stewart et al. 1972). Later airglow emissions were observed at the limb with SPICAM and IUVS spectrometers onboard MEX and MAVEN respectively (Bertaux et al. 2006; Leblanc et al. 2006; Jain et al. 2015; Gkouvelis et al. 2020). Ritter et al. (2019) investigated dayglow limb observations of oxygen FUV emission multiplets at 130.4 nm and 135.6 nm. These data were collected between 2014 and 2018 from IUVS instrument onboard MAVEN spacecraft. Recently NOMAD/UVIS instrument onboard EXOMars have provided dayglow observations of oxygen 557.7 nm and 630 nm emissions (Gérard, et al. 2020, 2021). Emirates Mars Ultraviolet Spectrometer (EMUS) on EMM is a first instrument, which is continuously measuring Extreme Ultraviolet and Far Ultraviolet (EUV and FUV) dayglow emissions. These emissions provide information about the atmospheric composition (both neutral and ions) and structure and can be used to study energy deposition, dynamics and chemistry. In Fig. 16.9 the green and red dots represent UVIS measurements corresponding to oxygen dayglow emissions of 557.7 nm and 630 nm respectively. The model calculations of these measurements were performed by Gérard, et al. (2020, 2021) at aphelion (Ls ~ 60°) (solid line) and perihelion (dashed line) (Ls ~ 251°) conditions for two different SZA 30° and 70°. The emission profile of oxygen 630 nm observed broad peak intensity 4.8 kR near 150 km in the Martian dayglow. The two peaks were observed for 557.7 nm emissions at ‘80 km and ~ 110 km with intensities 240 kR and 120 kR respectively.

16.5 Nightglow on Mars

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Fig. 16.9 Model calculations of the [O] 630 nm and [O] 557.7 dayglow limb intensities for two different SZA 30° and 70°. The UVIS observations of [O] 557.7 nm and [O] 630 nm lines are shown by the green and red dots respectively (from Gérard et al. 2021)

16.5 Nightglow on Mars The first detection of NO UV airglow on Mars nightside was reported by Bertaux et al., (2005). Later IUVS instrument onboard MAVEN also observed NO emission in the nighttime atmosphere of Mars (Stiepen et al. 2015, 2017; Schneider et al. 2020). This instrument observed 5 of the γ bands and 8 of the δ bands of the NO nightglow. The Fig. 16.10a, b shows NO nightglow brightness of γ and δ bands during orbit # 3241 on 30 May 2016 at Ls = 161° in MY33. In this figure the (a) image shows the enhancement near southern winter pole and also patchy equatorial enhancements of comparable magnitude. The figure (b) represents MUV spectral fit of the NO nightglow. The black line shows the spectrum observed, while the red line is fitted to the observed spectrum. The MUV channel spans between 185 to 299 nm corresponding to 174 spectral bins. The wavelength range between 195 to 221 nm for the 40 spectral bins recorded three strong delta bands (Δv' = 1, 2, 3) in Fig. 16.10b. The count rate of the NO nightglow emissions are shown on the y axis. This count rate is converted into the brightness of NO nightglow emission intensity, which is observed to be 1.68 ± 0.04 kR (Schneider et al. 2020). Several other measurements for nighglow emissions have been carried out in the Martian atmosphere. Haider and Oyama (2002) estimated nightglow limb intensities ~ 60 R and ~ 10 R at wavelengths 557.7 nm and 630 nm respectively. The model results were compared with the upper limit set by Mars 5 spectroscopic observations between wavelength region 300–800 nm (e.g. Krasnopolsky and Krysko 1976). After about three decades SPICAM (Fedorova et al. 2012), OMEGA (Bertaux et al. 2012) and Compact Reconnaissance Imaging Spectrometer (CRISM) (Clancy et al. 2010)

Fig. 16.10 a the image shows the expected enhancement near the southern pole, but also patchy equatorial enhancements of comparable magnitude. b MUV spectral fit of NO nightglow with black line observations and red line fitting to the observations (from Schneider et al. 2020)

146 16 Aurora and Airglow on Mars

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spectrometers onboard MEX and MRO observed the O2 (a1 Δg) nightglow at 1.27 μm on Mars. These instruments observed nightglow emission intensity ~ 5–15 MR at altitude 40–50 km for O2 (a1 Δg) at 1.27 μm (Krasnopolsky 2013; Clancy et al. 2017; Brecht and Kahre 2017). The infrared spectrometer onboard EXOMars/TGO is capable to perform emission intensity of O2 (a1 Δg) nightglow at 1.27 μm on Mars. This emission intensity of this band is not detected till now from TGO in the nightside atmosphere of Mars.

References Brecht, A.S., Kahre, M.: Nasa ANES Mars GCM with photochemistry: Modeling O2 IR nightglow emission. In: The Mars Atmosphere: Modelling and Observation, p. 4206 (2017) Bertaux, J.L., Gondet, B., Lefèvre, F., et al.: First detection of O2 1.27 μm nightglow emission at Mars with OMEGA/MEX and comparison with general circulation model predictions. J. Geophys. Res.: Planets 117(E11) (2012) Barth, C.A., Stewart, A.I., Hord, C.W., Lane, A.L.: Mariner 9 ultraviolet spectrometer experiment: Mars airglow spectroscopy and variations in Lyman alpha. Icarus 1; 17(2):457–468 (1972) Barth, C.A., Hord, C.W.: Mariner ultraviolet spectrometer: topography and polar cap: ultraviolet measurements reveal the topography of Mars and show that ozone may be adsorbed on the polar cap. Science 173(3993), 197–201 (1971) Bertaux, J.L., Leblanc, F., Witasse, O., et al.: Discovery of an aurora on Mars. Nature 435(7043), 790–794 (2005) Bertaux, J.L., Korablev, O., Perrier, S., et al.: SPICAM on Mars express: observing modes and overview of UV spectrometer data and scientific results. J. Geophys. Res.: Planets 111(E10) (2006) Chaufray, J.Y., Modolo, R., Leblanc, F., et al.: Mars solar wind interaction: formation of the Martian corona and atmospheric loss to space. J. Geophys. Res.: Planets 112(E9) (2007) Clancy, R.T., Wolff, M.J., Whitney, B.A., et al.: Extension of atmospheric dust loading to high altitudes during the 2001 Mars dust storm: MGS TES limb observations. Icarus 207(1), 98–109 (2010) Clancy, R.T., Smith, M.D., Lefèvre, F., et al.: Vertical profiles of Mars 1.27 μm O2 dayglow from MRO CRISM limb spectra: Seasonal/global behaviors, comparisons to LMDGCM simulations, and a global definition for Mars water vapor profiles. Icarus 293, 132–156 (2017) Clarke, J.T., et al.: Jupiter’s aurora. Jupiter: The planet, satellites and magnetosphere 1, 639–670 (2004) Deighan, J., Jain, S.K., Chaffin, M.S., et al.: Discovery of a proton aurora at Mars. Nat. Astron. 2(10), 802–807 (2018) Fedorova, A.A., Lefèvre, F., Guslyakova, S., et al.: The O2 nightglow in the Martian atmosphere by SPICAM onboard of Mars-Express. Icarus 1; 219(2), 596–608 (2012) Gérard, J.C., Hubert, B., Ritter, B., et al.: Lyman-α emission in the Martian proton aurora: Line profile and role of horizontal induced magnetic field. Icarus 321, 266–271 (2019) Gérard, J.C., Aoki, S., Willame, Y., et al.: Detection of green line emission in the dayside atmosphere of Mars from NOMAD-TGO observations. Nature Astronomy 4(11), 1049–1052 (2020) Gérard, J.C., Aoki, S., Gkouvelis, L., et al.: First observation of the oxygen 630 nm emission in the Martian dayglow. Geophys. Res. Lett. 28; 48(8), e2020GL092334 (2021) Gkouvelis, L., Gérard, J.C., Ritter, B., et al.: Airglow remote sensing of the seasonal variation of the Martian upper atmosphere: MAVEN limb observations and model comparison. Icarus 1(341), 113666 (2020)

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Gunell, H., Brinkfeldt, K., Holmström, M., et al.: First ENA observations at Mars: charge exchange ENAs produced in the magnetosheath. Icarus 182(2):431–438 (2006) Haider, S.A., Oyama, K.I.: Calculated electron flux and densities at 10–1000 eV in the dayside Martian ionosphere: comparison with MGS and Viking results (2002) Haider, S.A., Mahajan, K.K., Bougher, S.W., Schneider, N.M., et al.: Observations and modeling of Martian auroras. Space Sci. Rev. 218(4), 1–53 (2022) Halekas, J.S., Lillis, R.J., Mitchell, D.L.: MAVEN observations of solar wind hydrogen deposition in the atmosphere of Mars. Geophys. Res. Lett. 16; 42(21), 8901–8909 (2015a) Halekas, J.S., Taylor, E.R., Dalton, G., et al.: The solar wind ion analyzer for MAVEN. Space Sci. Rev. 195(1), 125–151 (2015b) Hughes, A., Chaffin, M., Mierkiewicz, E., et al.: Proton aurora on Mars: A dayside phenomenon pervasive in southern summer. J. Geophys. Res. Space. Phys. 124(12), 10533–10548 (2019) Jain, S.K., Stewart, A.I., Schneider, N.M., et al.: The structure and variability of Mars upper atmosphere as seen in MAVEN/IUVS dayglow observations. Geophys. Res. Lett. 16; 42(21), 9023–9030 (2015) Kallio, E., Barabash, S.: Atmospheric effects of precipitating energetic hydrogen atoms on the Martian atmosphere. J. Geophys. Res. Space. Phys. 106(A1), 165–177 (2001) Kallio, E., Janhunen, P.: Atmospheric effects of proton precipitation in the Martian atmosphere and its connection to the Mars-solar wind interaction. J. Geophys. Res. Space Phys. 106(A4), 5617–5634 (2001) Krasnopolsky, V.A., Krysko, A.A.: On the night airglow of the Martian atmosphere. Space Res. XVI, 1005–1008 (1976) Krasnopolsky, V.A.: Night and day airglow of oxygen at 1.27 μm on Mars. Planet Space Sci. 85, 243–249 (2013) Leblanc, F., Chaufray, J.Y., Lilensten, J., et al.: Martian dayglow as seen by the SPICAM UV spectrograph on Mars Express. J. Geophys. Res.: Planets 111(E9) (2006) Lundin, R., Winningham, D., Barabash, S., et al.: Plasma acceleration above Martian magnetic anomalies. Science 311(5763), 980–983 (2006) Mitchell, D.L., Lillis, R.J., Lin, R.P., et al.: A global map of Mars’ crustal magnetic field based on electron reflectometry. J. Geophys. Res. Planets 112(E1) (2007) Ritter, B., Gérard, J.C., Hubert, B.: Observations of the proton aurora on Mars with SPICAM on board Mars Express. Geophys. Res. Lett. 45(2), 612–619 (2018) Ritter, B., Gérard, J.C., Gkouvelis, L. et al.: Characteristics of Mars UV dayglow emissions from atomic oxygen at 130.4 and 135.6 nm: MAVEN/IUVS limb observations and modeling. J. Geophys. Res. Space Phys. 124(6), 4809–4832 (2019) Schneider, N.M., Deighan, J.I., Jain, S.K., et al.: Discovery of diffuse aurora on Mars Science. 350(6261), aad0313 (2015) Schneider, N.M., Jain, S.K., Deighan, J., et al.: Global aurora on Mars during the September 2017 space weather event. Geophys. Res. Lett. 45(15), 7391–7398 (2018) Schneider, N.M., Milby, Z., Jain, S.K., et al.: Imaging of Martian circulation patterns and atmospheric tides through MAVEN/IUVS nightglow observations. J. Geophys. Res. Space. Phys. 125(8), e2019JA027318 (2020) Soret, L., Gérard, J.C., Schneider, N., et al.: Discrete aurora on Mars: spectral properties, vertical profiles, and electron energies. J. Geophys. Res. Space Phys. 126(10), e2021JA029495 (2021) Stewart, A.I.: Mariner 6 and 7 ultraviolet spectrometer experiment: Implications of CO2 + , CO and O airglow. J. Geophys. Res. 77(1), 54–68 (1972)

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Stewart, A.I., Barth, C.A., Hord, C.W., Lane, A.L.: Mariner 9 ultraviolet spectrometer experiment: structure of Mars’ upper atmosphere. Icarus 17(2), 469–474 (1972) Stiepen, A., Gérard, J.C., Gagné, M.È., et al.: (2015) Ten years of Martian nitric oxide nightglow observations. Geophys. Res. Lett. 16; 42(3), 720–725 (2015) Stiepen, A., Jain, S.K., Schneider, N.M., et al.: Nitric oxide nightglow and Martian mesospheric circulation from MAVEN/IUVS observations and LMD-MGCM predictions. J. Geophys. Res. Space. Phys. 122(5), 5782–5797 (2017)

Chapter 17

Middle Ionosphere of Mars

Abstract The middle ionosphere of Mars is formed mainly due to ablation of meteors and charge exchange of metallic ions between altitudes 60 and 110 km (Withers et al., J Geophys Res Space Phys 113(A12), 2008). In general, the daytime ionosphere of Mars has been observed from RO experiment as E and F layers with electron densities that peaks at altitudes 110 and 130 km respectively (Kliore et al., Science 149:1243–1248, 1965; Rasool and Stewart, J Atmos Sci 28:869–878, 1971; Bougher et al., Geophys Res Lett 28:3091–3094, 2001; Thirupathaiah et al., Icarus 330:60–74, 2019). The RO experiment onboard MGS and MEX have also observed a third ionospheric layer in some of the electron density profiles at altitudes between 65 and 110 km (Pätzold et al., Science 310:837–839, 2005; Withers et al., J Geophys Res Space Phys 113(A12), 2008; Haider et al., J Geophys Res Space Phys 118:6786– 6794, 2013; Pandya and Haider, J Geophys Res Space Phys 119:9228–9245, 2014). The origin of this layer in the middle ionosphere has been attributed to ablation of meteors and charge exchange of Mg and Fe. The meteoroids are the major population in the solar system minor bodies at a distance of 1 AU from the Sun. It can interact with the atmosphere and reach on the surface of Earth, known as meteorites. While entering the Earth’s atmosphere, meteoroids experience several physical and chemical processes: collisions with atmospheric atoms, evaporation, ablation, temperature variation and decomposition. Various numerical models have been developed for the study of these processes of meteoroid impact ionizations (Briani et al. 2013). A significant proportion of meteoroids impacting the Earth’s atmosphere will be lost through the melting and evaporation occurring during atmospheric entry. Meteor ablation deposition will differ in several ways between Venus and Mars. Zook and Berg (1975) predicted that the flux of meteorites may be 20% larger in the orbit of Venus than in the orbit of Earth at 1 AU. The general features of the size and mass spectrum will remain the same, but impact velocities will be greater owing to the greater orbital velocity of both Venus and the meteor materials because of their proximity to the sun. Gadsden (1967) calculated meteor ablation in the Earth’s atmosphere, which showed the deposition peak at about 90 km. He also suggested that meteor peak density on Venus should be at 120 km. Keywords Meteoric layer · Stratosphere · Metalic ions

© Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_17

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17.1 MGS Measurements and Meteoric Layer The closest approach of comet 4015 Wilson-Harrington with Mars is shown in Fig. 17.1 on 23 May, 2005. In this figure the location of comet and positions of Earth, Mercury, Venus and Mars were simulated through the website http://ssd.jpl. nasa.gov/sbdb.cgi This comet left the debris and dust particles into the Martian atmosphere, which ablaze and produce a third broad peak at lowest altitude ~ 102 km. In Fig. 17.2 the electron density profile is observed by MGS, when comet 4015 WilsonHarrington crossed the Mars’ orbit on 23 May, 2005. In this figure the F peak at ~ 130 km is formed due to impact of solar EUV radiation (Ma et al. 2004; Haider et al. 2009; Mendillo et al. 2011; Withers et al. 2015; Thirupathaiah et al. 2019;). In this figure E-peak is not visible clearly. The Ls, latitude, longitude and SZA are given in this profile. The peak electron density during the meteor impact was observed to be ~ 1.85 × 1010 m−3. at 102 km. In presence of meteoric ablation, the electron density increased by about an order of magnitude between altitudes 60 and 90 km (Molina-Cuberos et al. 2003; Pandya and Haider 2012, 2014). The meteoric plasma layer is not a permanenr peak in the lower ionosphere of Mars. Haider et al. (2009) reported D, E, and F layers with concentrations 7 × 101 , 2.4 × 104 and 8.4 × 104 cm−3 at altitude range ~ 25–35, 100–112 and 125–135 km respectively. The meteoric ablation produced fourth ionization peak 1.85 × 104 cm−3 between D and E layers at altitude range ~ 80–105 km The electron density of this peak is lower than that of E and F peaks are produced by solar X-ray and EUV radiations respectively. Fig. 17.1 Comet 4015 Wilson-Harrington crossing Mars orbit from a closest distance 1.17 AU on 23 May, 2005 (from Pandya and Haider 2012)

17.3 Meteoric and Atmospheric Ions

153

Fig. 17.2 Altitude profiles of electron density measured by radio occultation experiment onboard MGS on 23 May, 2005 (from Pandya and Haider 2012)

17.2 MEX Measurements and Meteoric Layer In Fig. 17.3 the closest approach of comet P/2003WC7 (LINEAR Catlina) with Mars is shown on 18 April, 2004. In this figure the location of comet and positions of Earth, Venus, Mercury and Mars were simulated through the website http://ssd.jpl.nasa.gov/ sbdb.cgi. This comet also left the dust particles into the Martian atmosphere, which ablaze and produce a broad peak electron density 2 × 104 cm−3 at 90 km (Pätzold et al. 2005; Pandya and Haider 2014). In Fig. 17.4 this electron density profile is observed by MEX on 18 April, 2004, when comet P/2003WC7 was crossing the Mars orbit. Before and after this event the meteoric plasma layer was not detected. In this profile F peak at ~ 142 km is formed due to impact of solar EUV radiation (Pandya and Haider 2014). E-peak is not visible.

17.3 Meteoric and Atmospheric Ions An understanding of the complete daytime ionosphere measured by MGS/MEX can only be developed by using a theoretical model. Recently Haider et al. (2013) developed a coupled chemical model which produces three plasma layers in the

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Fig. 17.3 A simulated view of comet position with orbital path of Mercury, Earth, Venus and Mars. Comet P/2003 WC7 (LINEAR-Catalina) is at close approach to Mars on 18 April, 2004 (from Pandya and Haider 2014)

Fig. 17.4 The electron density profiles observed on 18 April 2004 by radio occultation experiment on board MEX at SZA 85° (from Pandya and Haider 2014)

17.5 Meteors at Mars Due to Encounter of Comet C/2013 A1

155

nighttime ionosphere of Mars at altitude ~ 80–100 km, ~ 120 km and ~ 160 km due to impact of the meteoroids, solar wind proton and electron respectively. The model results were compared with the nighttime RO instruments onboard MEX, Mars 4 and Mars 5 spacecrafts. This model is extended for the calculation of ion and electron densities in the daytime ionosphere of Mars (Pandya and Haider 2014). The densities of 21 ions (CO2 + , O2 + , CO+ , O+ , NO+ , N2 + , Mg+ , Fe+ , Si+ , MgO+ , FeO+ , SiO+ , MgCO2 + , MgO2 + , FeCO2 + , FeO2 + , SiCO2 + , SiO2 + , MgN2 + , FeN2 + , and SiN2 + ) and 10 neutral species (Mg, Fe, Si, MgO, FeO, SiO, MgCO3 , FeCO3 , MgO2 and FeO2 ) have been computed self-consistently for two specific days on 18 April, 2004 and 11 May, 2005 at SZA 85° and 82° respectively. The electron density is calculated using the charge neutrality under steady-state condition. This model also produces first and second peaks but at altitude range ~ 135–140 km and ~ 100–115 km due to impact of solar EUV and X-ray radiation respectively. The third peak is produced at altitudes 75 and 85 km due to ablation of meteoroids of masses 1 × 10–3 and 4.5 × 10–7 g respectively.

17.4 Meteoroid Flux The meteoroid flux is observed in the orbit of Earth. Therefore φe is scaled to φm at distance Dm = 1.5 φe from the sun as given below: φm = φe (Dm )0.5

(17.1)

where φe and φm are meteoroid fluxes on Earth and Mars respectively. The meteoroid flux is varying with mass of meteoroids. Figure 17.5 represents the meteoroid flux of different masses. These fluxes are reported for Earth’s atmosphere. The meteoroids collide with the atmospheric particles, ablaze and produce luminous phenomenon known as meteor showers. The meteors of 11.1, 30 and 72 km/s have been chosen in the Earth’s atmosphere for modeling of flux and meteoroid sputtering (Adolfson et al. 1996). Haider et al. (2013) have taken average velocity ~ 18 km/s for meteoroids of flux ~ 3.0 × 10–17 , 4.0 × 10–15 and 2.0 × 10–13 cm−2 s−1 with masses 1 × 10–2 , 3 × 10–4 and 1 × 10–5 gm respectively.

17.5 Meteors at Mars Due to Encounter of Comet C/2013 A1 The comet C/2013 A1 crossed the Mars orbit at minimum distance ~ 134,000 km on 19 October, 2014 at local time 18:35 UT (Farnocchia et al. 2014; Haider and Pandya 2015). The closest approach of this comet with Mars is shown in Fig. 17.6. In this figure the location of comet and positions of Earth, Venus, Mercury and Mars were

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17 Middle Ionosphere of Mars 10-5 Estimated Flux on Mars Grun Flux on Earth Zook Flux on Earth Le Sergeant and Lamy Flux on Earth

10-6

-1

10-9

Meteoroid Flux (m s )

10-8

-2

10-7

10-10 10-11 10-12 10-13 10-14 10-15 10-16 10-10

10-9

10-8

10-7

10-6

10-5

10-4

10-3

10-2

10-1

100

101

102

Mass (gm)

Fig. 17.5 Comparision of meteoroid flux with mass Estimated (dark triangle), Modeled by Grun et al. (1985) (red circles), Modeled by Zook et al. (1970), and Modeled by Le sergeant and Lamy (1980) on Earth

simulated through the website http://ssd.jpl.nasa.gov/sbdb.cgi. The NGIMS onboard MAVEN measured the plasma densities in the Martian environment (Benna et al. 2015). Figure 17.7a, b show dust count rate before and after the encounter of comet C/2013 A1 with Mars. This spectrum represents the presence of metallic ions Na+ , Mg+ , Fe+ , Al+ , K+ , Mn+ , Cr+ , Ni+ , Cu+ and Zn+ with possible rise of spectral lines for Si+ and Ca+ .

17.6 Comet C/2013 A1 Observations for Meteoric Ions The interplanetary dust and cometary dust grains entering the Martian atmosphere produce bright luminosity are known as meteors. Some meteors produce bright luminosity to exceed the brightness of the planets with visual magnitude ~ 10–4 gm are called fire-balls (Mann 2010). The optical meteors are observed with mass ranges from 10–5 < m < 100 gm. The meteors with mass < 10–2 gm are radio meteors. The dust particles with mass m > 10–5 are known as meteoroids (Mason 1971). Dust grains even lesser in mass ranging from 10–9 < m < 10–5 gm are widely affecting ionosphere known as micrometeoroids (Molina-Cuberos et al. 2003; Pandya and Haider 2015). Massive dust particles with mass 1 gm < m < 6 × 104 kg are known as meteorites, which burn in the atmosphere and also reach at the surface of the planet. Dust particles in space with mass ranging 10–12 to 10–15 gm which called β-meteoroids (Mann et al. 2014).

17.6 Comet C/2013 A1 Observations for Meteoric Ions

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Fig. 17.6 Comet C/2013 A1 encounter with Mars at 134,000 km (from Haider and Pandya 2015) Fig. 17.7 Mass spectra of dust counts a before and b after encounter of comet C/ 2013 A1 with Mars (from Benna et al. 2015)

Cosmic dust meteoroids are icy stones and contain uniform metallic compositions. The elemental abundances of the metallic constituents are Mg 0.08, Fe 0.05, Na 0.05, S 0.51, Ar 0.10, Al 0.08, Ca 0.06, Ni 0.04, Cr 0.01, P 0.01, K 0.003, Ti 0.002, Co 0.002, with relative to Si 1.0 (Sida 1969; Mason 1971; Fegle and Cameron 1987; Lodders and Fegley 1998; Jones 1997 and Janches et al. 2014). Sears and Dodd (1988) have estimated the major metallic constituents in the meteoritic abundances as Mg 14.4%, Si 13.6%, Fe 12.1%, Al 1.2%, Ca 0.82% and Na 0.80%. Later Jones (1997) reported that the abundant proportion of the atoms to the total atoms in cometary meteoroids is Mg (8.2%), Fe (5.9%), Si (24.2%) and O (61.7%).

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17.7 Chemistry of Meteoric Ions The ion-neutral collisions, electron-neutral collisions, dissociative recombination of molecular ions, ion-ion and electron–ion metal recombination, charge exchange and meteoroid impact ionization processes are used in the nighttime and daytime calculations of ion and electron densities between 50 and 200 km (Haider et al. 2013; Pandya and Haider 2014). These chemical reactions are used in Eq. (10.2) under charge neutrality steady state conditions. The chemical reactions for meteoric chemistry are given in Appendix 8, where Tn and Te are neutral and electron temperatures respectively. These temperatures are taken from Fox (1993).

17.8 Meteoric Model In Chap. 10, the coupled equation of ablation, motion and energy are solved [see Eqs. (10.35), (10.47), and (10.49)] using Adams-Bashferth method for the calculation of ion production rates due to impact of meteoroids in the Mars’ atmosphere. These equations are given below: cos θ · cos θ

[· A·ρ·V dV = 1 2 dz m /3 · δ /3

m 2/3 K 1 e(−K 2 /T ) Ʌs m 2/3 dm = −4 · A 2/3 · A − ρV 2 √ dz δ ·V 2Q δ 2/3 T cos θ

dT = dz

ρ V2A 1 2/ 2 c m / 3 δm 3

−

4 σ A · T4 1 2 c m /3δ /3 · V ( ) − K 2/ T ·e

(17.2)

(17.3)

(Ʌ − Ʌs ) −

4 A · K1 · Q √ 1/ 2 c · m 3 · δ /3 · V T

(17.4)

where V is the meteoroid velocity, A is the shape factor, δ is the meteoroid density, ρa is the atmospheric density, m is the mass of meteor substance, [ is drag coefficient, Ʌ is the heat transfer coefficient, ɅS is the sputtering coefficient, Q is the energy of evaporation, σ is the Stefan–Boltzman constant, θ is the entry angle, c is the heat capacity of the meteoroid substance, the parameters K1 and K2 define the dependence of the evaporation rate on temperature. The above parameters are given as Ʌ = [ = 1.0, θ = 45°, K1 = 6.92 × 1010 g / cm2 s, K2 = 5.78 × 104 degree, δ = 2.5 g/cm3 and T = 216 K (cf. Grun et al. 1985; Molina-Cuberos et al. 2003). The other parameters are given by Lebedinets et al. (1973) as A = 1.21, ɅS = 6 × 10–6 exp (T/290), Q = 6 × 1010 erg/g, c = 1 × 107 erg g−1 k−1 and σ = 5.68 × 10–5 erg cm−2 s−1 k− 4 . It is assumed that cometary meteoroids are composed of 61.7% oxygen, 24.2% silicon, 8.2% magnesium and

References

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5.9% iron in atoms of element to the total (Jones 1997). The major ions Fe+ , Mg+ and Si+ are produced due to meteor ablation in the nighttime and daytime ionosphere of Mars (Haider et al. 2013; Pandya and Haider 2014). The minimum velocities of Fe+ , Mg+ and Si+ are taken 9.4, 11.1 and 11 km/s respectively for the calculation of ionization probabilities.

References Adolfsson, L., Gustafson, B.A.S., Murray, C.D.: The Martian atmosphere as a meteoroid detector. Icarus 119, 144–152 (1996) Benna, M., Mahaffy, P.R., Grebowsky, J.M. et al.: Metallic ions in the upper atmosphere of Mars from the passage of comet C/2013 A1 (Siding Spring). Geophys. Res. Lett. 42 (2015). https:// doi.org/101002/2015GL064159 Bougher, S.W., Engel, S., Hinson, D.P., Forbes, J.M.: Mars Global Surveyor radio science electron density profiles: neutral atmosphere implications. Geophys. Res. Lett. 28(16), 3091–3094 (2001) Briani, G., et al.: Simulations of micrometeoroid interactions with the Earth atmosphere. Astron. Astrophys. 552(A53) (2013) Farnocchia, D., ChesleyS, R., Chodas, P.W., et al.: Trajectory analysis for the nucleus and dust of Comet C/2013 A1 (Siding Spring). Astrophys. J. 790, 114–121 (2014) Fegle, B., Jr., Cameron, A.G.W.: A vaporization model for iron/silicate fractionation in the Mercury protoplanet. Earth Planet. Sci. Lett. 82(3–4), 207–222 (1987) Fox, J.L., Brannon, J.F., Porter, H.S.: Upper limits to the nightside ionosphere of Mars. Geophys. Res. Lett. 20(13), 1339–1342 (1993) Gadsden, M.: Tables of meteor ablation. vol. 42. Institute for Telecommunication Sciences and Aeronomy (1967) Grun, E., Zook, H.A., Fechtig, H., Giese, R.H.: Collisional balance of the meteoritic comples. Icarus 62, 244–272 (1985) Haider, S.A., Pandya, B.M., Molina-Cuberos, G.J.: Nighttime ionosphere caused by meteoroid ablation and solar wind electron-proton-hydrogen impact on Mars: MEX observation and modeling. J. Geophys. Res. Space Phys. 118(10), 6786–6794 (2013) Haider, S.A., Pandya, B.M.: Probing of meteor showers at Mars during the encounter of comet C/ 2013 A1: predictions for the arrival of MAVEN/Mangalyaan. Geosci. Lett. 2, 8 (2015). https:// doi.org/101186/s40562-015-0023-2 Haider, S.A., et al.: D, E, and F layers in the daytime at high-latitude terminator ionosphere of Mars: comparison with Earth’s ionosphere using COSMIC data. J. Geophys. Res. Space Phys. 114(A3) (2009) Janches, D., Plane, J.M.C., Nesvorný, D., et al.: Radar detectability studies of slow and small zodiacal dust cloud particles I The case of Arecibo 430 MHz meteor head echo observations. Astrophys. J. 796(1), 41 (2014) Jones, W.: Theoretical and observational determinations of the ionization coefficient of meteors. Mon. Not. R. Astron. Soc. 288, 995–1003 (1997) Kliore, A., Cain, D.L., Levy, G.S.D., et al.: Occultation experiment: Results of the first direct measurement of Mars’s atmosphere and ionosphere. Science 149(3689), 1243–1248 (1965) Le Sergeant, D.L., Lamy, P.L.: On the size distribution and physical properties of interplanetary dust grains. Icarus 43(3), 350–372 (1980) Lodders, S.G., Fegley, J.B.: The Planetary Scientist’s Companion, p. 371. Oxford University Press, New York (1998) Ma, Y., et al.: Three-dimensional, multispecies, high spatial resolution MHD studies of the solar wind interaction with Mars. J. Geophys. Res. Space. Phys. 109(A7) (2004) Mann, I.: Interstellar dust in the solar system. Ann. Rev. Astron Astrophys. 48, 173–203 (2010)

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Mann, I., Meyer-Vernet, N., Czechowski, A.: Dust in the planetary system: dust interactions in space plasmas of the solar system. Phys. Rep. 536(1), 1–39 (2014). http://doi.org/101016/jph ysrep201311001 Mason, B.: Handbook of Elemental Abundances of the Elements in Meteorites, p. 1971. Gordon and Breach, Newark (1971) Mendillo, M., Lollo, A., Withers, P. et al.: Modeling Mars’ ionosphere with constraints from sameday observations by Mars Global Surveyor and Mars Express. J. Geophys. Res. Space Phys. 116(A11), (2011) Molina-Cuberos, G.J., Witass, O., Lebreton, J.P., et al.: Meteoric ions in the atmosphere of Mars. Planet. Space Sci. 51(3), 239–249 (2003) Pandya, B.M., Haider, S.A.: Meteor impact perturbation in the lower ionosphere of Mars: MGS observations. Planet. Space Sci. 63, 105–109 (2012) Pandya, B.M., Haider, S.A.: Numerical simulation of the effects of meteoroid ablation and solar EUV/X-ray radiation in the dayside ionosphere of Mars: MGS/MEX observations. J. Geophys. Res. Space. Phys. 119(11), 9228–9245 (2014) Pätzold, M., Tellmann, S., Hausler, B., et al.: A sporadic third layer in the ionosphere of Mars. Science 310(5749), 837–839 (2005) Rasool, S.I., Stewart, R.W.: Results and interpretation of the S-band occultation experiments on Mars and Venus. J. Atmos. Sci. 28(6), 869–878 (1971) Sears, D.W.G., Dodd, R.T.: Overview and classification of meteorites. In: Kerridge, J.F., Matthews, M.S. (eds.) Meteorites and the Early Solar System, pp. 3–31. University of Arizona Press, Tucson (1988) Sida, D.W.: The Production of Ionsand electrons by meteoric processes. Mon. Not. R. Astr. Soc. 143, 37–47 (1969) Thirupathaiah, P., Shah, S.Y., Haider, S.A.: Characteristics of solar X-ray flares and their effects on the ionosphere and human exploration to Mars: MGS radio science observations. Icarus 330, 60–74 (2019) Withers, P., Morgan, D.D., Gurnett, D.A.: Variations in peak electron densities in the ionosphere of Mars over a full solar cycle. Icarus 251, 5–11 (2015) Withers, P., Mendillo, M., Hinson, D.P., Cahoy, K.: Physical characteristics and occurrence rates of meteoric plasma layers detected in the Martian ionosphere by the Mars Global Surveyor Radio Science Experiment. J. Geophys. Res. Space Phys. 113(A12) (2008) Zook, H.A., Flaherty, R.E., Kessler, D.J.: Meteoroid impacts on the Gemini windows. Planet. Space Sci. 18, 953–964 (1970) Zook, H.A., Berg O.E.: A source for hyperbolic cosmic dust particles. Planet. Space Sci. 23(1), 183–203 (1975)

Chapter 18

Lower Atmosphere of Mars

Abstract The lower atmosphere of Mars exists below 100 km, where gases are mixed and eddy diffusion is dominant. It is characterized by a strong coupling between pressure, temperature, neutral density and winds. MGS and MEX have observed temperature, pressure and total density in the lower atmosphere of Mars with radio occultation experiment (Bougher et al. in Geophys Res Lett 28:3091–3094, 2001; Hinson et al. in J Geophys Res Planets 104:26997–27012, 1999; Pätzold et al. in Science 310:837–839, 2005). The neutral winds are mapped at 150 km by NGIMS onboard MAVEN to study the wind circulation and climatology of Mars (Benna et al. in Science 366:1363–1366, 2019). The troposphere occurs in the lower atmosphere between 0 and 50 km. The diurnal temperature variability is high at the surface due to low thermal inertia. During dust storm the suspended dust particles can reduce the surface diurnal temperature. There is no stratosphere in the lower atmosphere of Mars due to lack of ozone in the middle atmosphere. The Mars has mesosphere between 50 and 100 km above the troposphere. The lapse rates are lower in the mesosphere than the troposphere. Keywords Wind circulation · Neutral density · Pressure

18.1 Temperature and Pressure Measurements The seasonal variations of atmospheric temperature and pressure were measured in situ by Viking 1 and 2 at landing sites (Hess et al. 1976). Figure 18.1 shows the seasonal variability of the pressure observed by Viking 1 and 2 Landers on the surface of Mars. The first minimum of the pressure near Ls = 150 occurs during southern winter when a great part of the atmosphere is trapped in the southern polar cap. The second maximum near Ls = 430 corresponds to the northern winter much shorter and less cold than southern winter (Hourdin et al. 1993; Haider and Mahajan 2014). Vertical profiles of atmospheric density, pressure and temperature have also been observed in situ from two Mars’ rovers, Spirit and Opportunity, which landed in the afternoon at local times ~ 14:16 h and 13:13 h on 4 January and 25 January 2004 at coordinates: − 14.5° N, 175.5° E and − 1.9° N, 354.5° E respectively (Withers and © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_18

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Fig. 18.1 Time evolution of the surface pressure observed by Viking 1/2 Landers during first 2587 three Martian years of the mission (from Haider and Mahajan 2014)

Smith 2006). Using remote sensing method MGS and MEX observed temperature, pressure and total density from radio occultation experiment in northern and southern troposphere of Mars (Pätzold et al. 2005). These and earlier observations have shown that up to 45 km the temperatures are highly variable and this variability depends on latitude, season and dust content of the atmosphere (Haider et al. 2008; Tellmann et al. 2013; Haider and Mahajan 2014).The TES onboard MGS (cf. Smith 2004), THEMIS onboard Mars Odyssey (Grassi et al. 2005; Tellmann et al. 2013) and MCS onboard MRO (Hinson and Wilson 2004; Heavens et al. 2011) have also provided temperature profiles and their variability in the atmosphere of Mars.

18.2 Neutral Density and Composition The SAM instrument onboard Curiosity rover observed the compositions of the Mars’ atmosphere at the surface: CO2 (95%), Ar (1.9%), N2 (2.6%), O2 (0.16%) and CO (0.06%) (Stern et al. 2015). Very small amount of O3 (0.03%) (Khayat et al. 2021; Patel et al. 2021), H2 O (0.03%) (Jakosky and Farmer 1982), CH4 (0.05–60 ppb) (Moores et al. 2019), SO2 (~ 0.2 ppb) (Ojha et al. 2019) and NO (70–1100 ppm) (Stern et al. 2015) are also measured varying from parts per million (ppm) to parts per billion (ppb) near the surface of Mars. There is no direct measurement of density profiles in the lower atmosphere of Mars. The first direct measurement of the density profiles in the upper atmosphere of Mars was performed by Viking 1 and 2 Landers on 20 July 1976 and 3 September, 1976 at latitudes 22.5° N, 48° W and 48° N, 22° W respectively at solar zenith angle 44°. The neutral densities in the lower atmosphere of Mars have been estimated by

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several authors (Molina-Cuberos et al. 2002; Haider et al. 2011; Millour et al. 2014, 2017; Bougher et al. 2015, 2017). These profiles have been described in Sect. 20.5.

18.3 Dynamics of the Lower Atmosphere The dynamics in the lower atmosphere of Mars can be understood from the behavior of CO2 , water vapor and dust due to solar and seasonal variability (e.g. Barnes et al. 2017a, b). The thermal structure of the lower atmosphere depends greatly on the surface temperature, dust content and aerosol dynamics. It is also affected by season, latitude, orbital motion and difference between Mars’ perihelion and aphelion distance from the Sun. The standard Martian seasons mainly drive changes in surface and atmospheric temperatures at high latitudes. The orbital seasons have large impact on the surface and the atmospheric temperatures at low latitudes (Almatroushi et al. 2021). The dust also plays an important role in the atmospheric dynamics. It is an abundant constituent on the surface and in the atmosphere where it resides mainly in the lower-middle atmosphere. It is an absorber for solar radiation and emitter/absorber for infrared radiation, which are strongly affecting the thermal structure of the Martian atmosphere (Kahre et al. 2017; Wolff et al. 2017). The seasonal variation of dust can be divided into two main seasons: (1) non-dusty season (Ls ~ 0°–135°) during northern spring and northern summer, where column dust opacity is low, and (2) dusty season (Ls ~ 135°–360°) during southern spring and southern summer where column dust opacity is high (Smith 2019; Kahre et al. 2017; Kass et al. 2016). The other important constituent in the lower atmosphere of Mars is water vapor, which is observed using absorption bands in the thermal IR and near IR (Smith et al. 2018; Montmessin et al. 2017). The global average of water column abundance is ~ 10 pr-µm. The seasonal cap leads to peak values up to ~ 50 pr-µm during northern summer. The water vapor condenses for the production of water ice clouds. The H2 O variability can play a major role in the radiative budget of the lower atmosphere (e.g. Clancy et al. 1996; Madeleine et al. 2012).The ozone is anti-correlated with water vapor. The water vapor photodissociates in the atmosphere and increases the abundance of odd hydrogen, which destroys ozone (Haider et al. 2019). The ozone is observed in the lower and middle atmosphere through the absorption of Hartley band at 255 nm (Clancy et al. 2016; Haider et al. 2022).

18.4 Atmospheric Wind Because the atmosphere of Mars is thin, high wind velocities are needed to move sand and dust from the surface of Mars. The Viking Landers measured speeds up to 113 km/h during dust storms (Pollack et al. 1979; Kuroda et al. 2008). Recently the measurements of winds are mapped at altitude ~ 150 km by NGIMS onboard MAVEN

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(Benna et al. 2019). In Fig. 18.2 the average neutral wind speeds and flow directions observed by MAVEN with respect to local time and latitudes were shown. These neutral winds are measured during 33 monthly campaigns between April, 2016 and December, 2018. The season of each campaign is indicated by Ls. A neutral wind flow map generated by Mars Global Ionosphere Thermosphere Model (MGITM) is also shown by grey arrows for comparison. The magnitude of the orbit-to-orbit variability of the winds is parameterized by the Generalized Coefficient of Variation (GCV), a consolidated measure of changes in both flow direction and magnitude. The details about the winds circulation during campaigns C# 1 to C # 33 are provided by Benna et al. (2019) and Roeten et al. (2019). In Fig. 18.3 the zonal, meridional and vertical winds represent seasonal variability due to formation of polar ice caps on Mars. Below 10 km altitude, the zonal wind speed increased up to ~ 22 m/s in MY28 and 15 m/s in MY 29 at Ls ~ 270o (left panel) (Haider et al. 2022). Above 20 km altitude, the effect of dust storm is almost negligible on zonal winds. The meridional wind velocities are not affected by the dust storm below 10 km (middle panel). The vertical winds in the tropics are mostly in the upward direction at all altitudes (right panel).

18.5 Modeling of the Lower Atmosphere The photochemical models have predicted concentrations of O2 , O3 , CO2 , CO and other neutral species in the lower atmosphere of Mars (Belton and Hunten 1966; Parkinson and Hunten 1972; Rodrigo et al. 1990; Krasnopolsky 2003). MolinaCuberos et al. (2002), Haider et al. (2011) and Haider and Mahajan (2014) have reported neutral model atmosphere of 12 gases (CO2 , N2 , Ar, O2 , CO, H2 , H2 O, O, O3 , NO, NO2 , and HNO3 ) in the lower atmosphere of Mars. The density profiles of these gases are plotted in Fig. 18.4. Mars Climate Database (MCD) provides meteorological parameters like air density, temperature, wind and mixing ratios using General Circulation Model (GCM) at different altitude, latitude and longitude in the Martian atmosphere (Millour et al. 2014) (website http://www-mars.lmd.jussieu.fr). This model has been developed for low, medium and high dust storm conditions. The solar EUV, X-rays and particle radiations are highly variable and therefore the atmosphere of Mars also changes in response to these variations. The MCD model includes the effect of such variations. The climatology scenario is also provided in the model for minimum, medium and maximum solar activity conditions. The MGITM also provides fundamental physical parameters like wind, gravity/ planetary waves, thermal structure, compositional and dynamical structures of Mars atmosphere from ground to the exosphere (0–250 km) (Bougher et al. 2015). In this model lower, middle and upper atmospheric processes are included and it is based on the formulations used in MTGCM (Bougher et al. 2017) and the legacy of NASA Ames Mars Global Circulation Model (MGCM) (e.g. Haberle et al. 1999).

Fig. 18.2 Dominant neutral winds observed using NGIMS during 33 monthly campaigns from April 2016 to December 2018 (from Benna et al. 2019)

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Fig. 18.4 Neutral model atmosphere in the lower atmosphere of Mars (taken from Molina-Cuberos et al. 2002; Haider et al. 2011);

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Haider, S.A., Pabari, J.P., Masoom, J., Shah, S.Y.: Schumann resonance frequency and conductivity in the nighttime ionosphere of Mars: a source for lightning. Adv. Space Res. 63(7), 2260–2266 (2019) Haider, S.A., Sheel, V., Singh, V., Maguire, W.C., Molina-Cuberos, G.J.: Model calculation of production rates, ion and electron densities in the evening troposphere of Mars at latitudes 67 N and 62 S: Seasonal variability. J. Geophys. Res. Space Phys. 113(A8) (2008) Haider, S.A., Mahajan, K.K., Kallio, E.: Mars ionosphere: a review of experimental results and modeling studies. Rev. Geophys. 49(4) (2011) Heavens, N.G., Richardson, M.I., Kleinböhl, A. et al.: Vertical distribution of dust in the Martian atmosphere during northern spring and summer: high-altitude tropical dust maximum at northern summer solstice. J. Geophys. Res. Planets. 116(E1) (2011) Hess, S.L., Henry, R.M., Leovy, C.B., et al.: Mars climatology from Viking 1 after 20 sols. Science 194(4260), 78–81 (1976) Hinson, D.P., Simpson, R.A., Twicken, J.D., et al.: Initial results from radio occultation measurements with Mars Global Surveyor. J. Geophys. Res. Planets 104(E11), 26997–27012 (1999) Hinson, D.P., Wilson, R.J.: Temperature inversions, thermal tides, and water ice clouds in the Martian tropics. J. Geophys. Res. Planets 109(E1) (2004) Hourdin, F., Van, P.L., Forget, F., Talagrand, O.: Meteorological variability and the annual surface pressure cycle on Mars. J. Atmos. Sci. 50(21), 11 (1993) Jakosky, B.M., Farmer, C.B.: The seasonal and global behavior of water vapor in the Mars atmosphere: complete global results of the Viking atmospheric water detector experiment. J. Geophys. Res. Solid Earth 87(B4), 2999–3019 (1982) Kahre, M.A., Murphy, J.R., Newman, C.E., et al.: The Mars dust cycle. Atmos. Clim. Mars. 18, 295 (2017) Kass, D.M., Kleinböhl, A., McCleese, D.J., et al.: Interannual similarity in the Martian atmosphere during the dust storm season. Geophys. Res. Lett. 43(12), 6111–6118 (2016) Khayat, A.S., Smith, M.D., Wolff, M. et al.: ExoMars TGO/NOMAD-UVIS vertical profiles of ozone: 2 the high-altitude layers of atmospheric ozone. J. Geophys. Res. Planets 126(11), e2021JE006834 (2021) Krasnopolsky, V.A.: Spectroscopic mapping of Mars CO mixing ratio: detection of north-south asymmetry. J. Geophys. Res. Planets 108(E2) (2003) Kuroda, T., Medvedev, A.S., Hartogh, P., Takahashi, M.: Semiannual oscillations in the atmosphere of Mars. Geophys. Res. Lett. 35(23) (2008) Madeleine, J.B., Forget, F., Millour, E., Navarro, T., Spiga, A.: The influence of radiatively active water ice clouds on the Martian climate. Geophys. Res. Lett, 39(23) (2012) Millour, E., et al.: The Mars Climate Database (MCD Version 53) EGU General Assembly Conference Abstracts (2017a) Millour, E., Forget, F., Spiga, A., et al.: A new Mars climate database v5 1. In: The Fifth International Workshop on the Mars Atmosphere: Modelling and Observations, p. id-1301 (2014) Molina-Cuberos, G.J., Lichtenegger, H., Schwingenschuh, K., et al.: Ion-neutral chemistry model of the lower ionosphere of Mars. J. Geophys. Res. Planets 107(E5), 3–1 (2002) Montmessin, F., Korablev, O., Lefèvre, F., Bertaux, J.L., Fedorova, A., Trokhimovskiy, A., Chapron, N.: SPICAM on Mars express: a 10 year in-depth survey of the Martian atmosphere. Icarus 297, 195–216 (2017) Moores, J.E., King, P.L., Smith, C.L., Martinez, G.M., Newman, C.E., Guzewich, S.D., Schuerger, A.C.: The methane diurnal variation and microseepage flux at Gale crater, Mars as constrained by the ExoMars trace gas orbiter and curiosity observations. Geophys. Res. Lett. 46(16), 9430–9438 (2019) Ojha, L., Karunatillake, S., Iacovino, K.: Atmospheric injection of sulfur from the Medusae Fossae forming events. Planet Space Sci. 179, 104734 (2019) Parkinson, T.D., Hunten, D.M.: Spectroscopy and acronomy of O2 on Mars. J. Atmos. Sci. 29(7), 1380–1390 (1972)

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Patel, M.R., Sellers, G., Mason, J.P., et al.: ExoMars TGO/NOMAD-UVIS vertical profiles of ozone: 1 seasonal variation and comparison to water. J/ Geophys. Res. Planets 126(11), e2021JE006837 (2021) Pätzold, M., Tellmann, S., Hausler, B., et al.: A sporadic third layer in the ionosphere of Mars. Science 310(5749), 837–839 (2005) Pollack, J.B., Colburn, D.S., Flasar, F.M.: Properties and effects of dust particles suspended in the Martian atmosphere. J. Geophys. Res. Solid Earth 84(B6), 2929–2945 (1979) Rodrigo, R., Garcia-Alvarez, E., Lopez-Gonzalez, M.J., Lopez-Moreno, J.J.: A nonsteady onedimensional theoretical model of Mars’ neutral atmospheric composition between 30 and 200 km. J. Geophys. Res. Solid Earth 95(B9), 14795–14810 (1990) Roeten, K.J., Bougher, S.W., Benna, M., et al.: MAVEN/NGIMS thermospheric neutral wind observations: interpretation using the M-GITM general circulation model. J. Geophys. Res. Planets 124(12), 3283–3303 (2019) Smith, M.D.: Interannual variability in TES atmospheric observations of Mars during 1999–2003. Icarus 167(1), 148–165 (2004) Smith, M.D.: THEMIS observations of the 2018 Mars global dust storm. J. Geophys. Res. Planets 124(11), 2929–2944 (2019) Smith, M.D., Daerden, F., Neary, L., Khayat, A.: The climatology of carbon monoxide and water vapor on Mars as observed by CRISM and modeled by the GEM-Mars general circulation model. Icarus 301, 117–131 (2018) Stern, J.C., Sutter, B., Freissinet, C., et al.: Evidence for indigenous nitrogen in sedimentary and aeolian deposits from the curiosity rover investigations at Gale crater. Mars Proc. Natl. Acad. Sci. 112(14), 4245–4250 (2015) Tellmann, S., Pätzold, M., Häusler, B., et al.: The structure of Mars lower atmosphere from Mars Express Radio Science (MaRS) occultation measurements. J. Geophys. Res. Planets 118(2), 306–320 (2013) Withers, P., Smith, M.D.: Atmospheric entry profiles from the Mars exploration rovers spirit and opportunity. Icarus 185(1), 133–142 (2006) Wolff, M.J., Lopéz-Valverde, M., Madeleine, J.B., et al.: Radiative process: techniques and applications. Atmos. Clim. Mars 18, 106 (2017)

Chapter 19

Trace Gases of Mars Atmosphere

Abstract The atmosphere of Mars is primarily composed of CO2 (~ 95%) with small amounts of N2 (~ 2.6%), Ar (~ 1.9%, O2 (0.16%) and CO (~ 0.06%). The O3 , H2 O, SO2 , NO and CH4 are the trace gases in the troposphere of Mars varying from parts per million to parts per billion (e.g. Khayat et al., J Geophys Res Planets 126(11):e2021JE006834, 2021; Farmer et al., J Geophys Res Space 82:4225–4248, 1977; Stern et al., Mars Proc Natl Acad Sci 112:4245–4250, 2015; Ojha et al., Planet Space Sci 179:104734, 2019). The CH4 , H2 O and O3 are involved in the green house effect, which keeps the planet warmer than it would be without an atmosphere. The SO2 is produced in the Martian atmosphere from natural phenomena such as volcanic eruptions. The NO may likely come from the energy released in meteorite impact. Keywords Trace gases · Measurements · Dynamics

19.1 Ozone Ozone is the most reactive species measured in the Martian atmosphere. It is formed from the product of O and O2 by three body reaction using CO2 as third body. The O3 is a sensitive tracer of the hydrogen photochemistry that stabilizes the Mars CO2 atmosphere. It also provides important information on the conditions of habitability of the planet, not only by determining the ultraviolet flux reaching the surface, but also as an indirect tracer of the oxidizing capacity of the atmosphere by hydrogen oxides. It was detected for the first time by absorptions of Hartley band in UV as well as IR features by the UV spectrometer on board Mariner 7 flyby (Barth and Hord 1971). There are only few instruments which explored ozone in Martian atmosphere. The UV spectrometer on board Mariner 9 (Barth et al. 1973), SPICAM on board MEX (Perrier et al. 2006), MARCI on board MRO (Clancy et al. 2016) and NOMAD on board TGO measured the ozone. Recently, vertical profiles of ozone observed by SPICAM have shown the presence of two distinct ozone layers at low to mid latitudes (Lebonnois et al. 2006). The first peak occurred near the at ~ 20 km. The second peak was detected at ~ 45 km. Montemessin and Leferve (2013) have © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_19

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Fig. 19.1 Vertical profiles of ozone over the north and south polar latitudes measured by NOMAD/ TGO (from Khayat et al. 2019)

reported a third layer of ozone over the winter south pole between 35 and 60 km. The NOMAD instrument also observed the altitudinal, latitudinal and seasonal distribution of ozone. (Khayat et al. 2021; Patel et 2021). Figure 19.1 shows the altitude profiles of ozone at Ls ~ 0–25 in the high northern (60° N–90° N) and southern (60° S–90° S) latitude regiom.

19.2 Water Vapour Water vapour exhibits the most important role in the Martian photochemistry. It produces HOx radicals, which stabilizes the atmosphere. First detailed knowledge of water vapour came from the MAWD instrument onboard Viking orbiter (Farmer et al. 1977; Jakosky and Farmer 1982). The MAWD measured water vapour column abundance as high as 100 pr-μm at Ls = 120° over northern polar region and a global and seasonal average water abundance of 10 pr-μm. The solar occultation observations on board Phobos 2 orbiter from the Auguste experiment yielded vertical distribution of water vapour in the altitude range 10–50 km (Krasnopolsky et al. 1991; Rodin et al. 1997). The vertical profiles of water vapour have been investigated by the solar occultation measurements with SPICAM onboard MEX (Maltagliati et al. 2013; Fedorova et al. 2018) and by the limb measurements with CRISM and MCS onboard MRO (Clancy et al. 2017). Recently, the solar occultation measurements were also carried out by the NOMAD instrument (Vandaele et al. 2018) during and

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Fig. 19.2 Altitude profiles of a H2 O and (b) HDO in presence and absence of GDS 2018 (from Vandaele et al. 2018)

after the GDS of MY34. This instrument observed altitude profiles of H2 O and HDO mixing ratios in the presence and absence of GDS.

19.3 Sulfur Dioxide If volcanoes have been active in recent Martian history, it would be expected to find SO2 in the current Martian atmosphere (Krasnopolsky 2012). The sulfur in the Martian atmosphere can form sulfate aerosols and play an important role in the global climate and in the atmospheric chemistry (Ojha et al. 2019). There have been reports of several observations that the Martian regolith is rich in sulfur (Berger et al. 2016). Remote sensing observations have found evidence for massive and widespread deposits of sulfates and bulk regolith throughout the planet and the Martian meteorites are known to contain sulfides and sulfates (Ding et al. 2015). The sulfur on the Martian surface has been derived from the mantle by various magnetic and hydrothermal processes (Franz et al. 2019). In situ sulfur isotope measurements from Curiosity rover resulted in wide ranges. The main source of sulfur on Mars is the volcanic exhalation of SO2 and H2 S gas (Ojha et al. 2019). The SO2 is photo-dissociated to produce sulfur monoxide (SO) in the atmosphere. The SO can be photo dissociated to form sulfur atoms. The reaction of SO with O2 or OH recycles sulfur atom back to SO2 while the reaction with O3 forms sulfur trioxide (SO3 ), which quickly combines

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with water vapour and produce H2 SO4 and subsequently condenses in the lower atmosphere.

19.4 Methane The presence of methane may enhance the habitability and also could be a signature of life (Giuranna et al. 2019). There are many reports of methane in Martian atmosphere with detection from orbit around Mars from terrestrial telescope and ground based measurements in Gale Crater (4.5° S, 137° E) on the curiosity rover (Etiope and Oehler 2019). In 2014, NASA has reported that Curiosity rover detected a tenfold increase in methane in the Martian atmosphere (Webster et al. 2015, 2018). The largest concentration of methane detected by the Curiosity rover shows a spike at 21 ppbv, during late June 2019 (Pla-Garcia et al. 2019). The methane is photodissociated into CH, 3 CH2 , 1 CH2 , and CH3 . The 3 CH2 reacts with the background gas CO2 to form CH2 O at altitudes above 80 km. This explains the upper peak in the CH2 O concentration at high altitudes. At low altitudes, the oxidation of methane by O (1 D) and OH forms CH3 . CH3 reacts with O to form CH2 O and with OH to form CH3 OH. Figure 19.3 represents the seasonal variability of CH4 concentration obtained from Mars Science Laboratory (MSL) during MY32, MY33, and MY34. In this figure upper limit of CH4 is also shown by open triangle with each measurements of MSL. The measurements of CH4 using ACS and NOMAD instruments onboard TGO are also shown. The upper limits of CH4 (95% confidence limit) obtained by TGO are compared to seasonally variable background methane measured by MSL during three Martian Years (MY32, MY33, and MY34). The colour scale gives the latitude of TGO sampling. The TGO is able to detect concentration about 10 times lower than 0.4 p.p.b.v. This difference can occur due to measurements of CH4 at two different locations. MSL measurements were obtained at the bottom of the Gale crater near the equator, whereas TGO measurements were achieved near polar latitudes at a few kilometres above the surface of Mars.

19.5 Nitric Oxide MSL has also detected nitrogen on the surface of Mars from release during heating of Martian sediments (Stern et al. 2015). The nitrogen was detected in the form of NO and could be released from the breakdown of nitrates during heating. There is no evidence to suggest that the fixed nitrogen molecules found by any form of NASA team were created the life. The source of Mars is in hospitable for known forms of life. NASA curiosity team thinks that the nitrates are ancient and likely came from non-biological processes like meteorite impact and lightning in Mars’ distant past. The Curiosity team has found evidence that other ingredients for life, such as liquid

19.5 Nitric Oxide

175

Fig. 19.3 TGO detection attempts combined with the MSL observations. Upper limits for CH4 (95% confidence limit) obtained by TGO (ACS and NOMAD) are compared to seasonally variable background methane as measured by MSL during three Martian years (MY32, MY33 and MY34). The colour scale gives the latitude of the TGO sampling (from Korablev et al. 2019).

water and organic maters, were present on Mars at the Curiosity site in Gale crater billions of years ago. Figure 19.4 represents the altitude profile of NO density obtained from Viking neutral mass spectrometer (red colour). The Imaging Ultraviolet Spectrograph (IUVS) onboard MAVEN also observed NOγ band between 213 and 225.5 nm. The IUVS NO number density at 117 km is 5 times smaller than those measured by Viking mass spectrometer over 40 years ago. IUVs observations are consistent with the photochemical model. Viking measurement of NO density is larger by an order of magnitude with the photochemical mode. The Viking and IUVS photochemical model are carried out by Fox (2004). These models were computed at SZA 44° and 75° for Viking and IUVS measurements respectively.

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Fig. 19.4 NO densities retrieved using the average IUVS (black line with symbols). The shaded area represents the 1-σ natural variability around the average limb radiance. The red line with symbols is the NO densities reported from the Viking 1 entry measurements (Nier and McElroy 1977). The black dashed curve is a photochemical model prediction for the IUVS measurements, and the red dashed curve is the same but for the Viking measurements (from Stevens et al. 2019).

References Barth, C.A., Hord, C.W.: Mariner ultraviolet spectrometer: topography and polar cap: ultraviolet measurements reveal the topography of Mars and show that ozone may be adsorbed on the polar cap. Science 173(3993), 197–201 (1971) Barth, C.A., Hord, C.W., Stewart, A.I., et al.: Mariner 9 ultraviolet spectrometer experiment: seasonal variation of ozone on Mars. Science 179(4075), 795–796 (1973) Berger, J.A., Schmidt, M.E., Gellert, R., et al.: A global Mars dust composition refined by the alpha-particle x-ray spectrometer in Gale Crater. Geophys. Res. Lett. 43(1), 67–75 (2016) Clancy, R.T., Smith, M.D., Lefèvre, F., et al.: Vertical profiles of Mars 1.27 μm O2 dayglow from MRO CRISM limb spectra: seasonal/global behaviors, comparisons to LMDGCM simulations, and a global definition for Mars water vapor profiles. Icarus 293, 132–156 (2017) Clancy, R.T., Wolff, M.J., Lefèvre, F., et al.: Daily global mapping of Mars ozone column abundances with MARCI UV band imaging. Icarus 266, 112–133 (2016) Ding, S., Dasgupta, R., Lee, C.T.A., Wadhwa, M.: New bulk sulfur measurements of Martian meteorites and modeling the fate of sulfur during melting and crystallization—implications for sulfur transfer from Martian mantle to crust–atmosphere system. Earth Planet. Sci. Lett. 409, 157–167 (2015) Etiope, G., Oehler, D.Z.: Methane spikes, background seasonality and non-detections on Mars: a geological perspective. Planet Space Sci. 168, 52–61 (2019) Farmer, C.B., Davies, D.W., Holland, A.L., et al.: Mars: water vapor observations from the Viking orbiters. J. Geophys. Res. Space 82(28), 4225–4248 (1977) Fedorova, A., Bertaux, J.L., Betsis, D., et al.: Water vapor in the middle atmosphere of Mars during the 2007 global dust storm. Icarus 300, 440–457 (2018) Fox, J.L.: Response of the Martian thermosphere/ionosphere to enhanced fluxes of solar soft X rays. J. Geophys. Res. Space. Phys. 109(A11), (2004) Franz, H.B., King, P.L., Gaillard, F.: Sulfur on Mars from the atmosphere to the core. In Volatiles in the Martian Crust, pp. 119–183. Elsevier (2019)

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Giuranna, M., Viscardy, S., Daerden, F., et al.: Independent confirmation of a methane spike on Mars and a source region east of Gale Crater. Nat. Geosci. 12(5), 326–332 (2019) Jakosky, B.M., Farmer, C.B.: The seasonal and global behavior of water vapor in the Mars atmosphere: complete global results of the Viking atmospheric water detector experiment. J. Geophys. Res. Solid Earth 87(B4), 2999–3019 (1982) Khayat, A.S.J., Smith, M.D., Guzewich, S.D.: Understanding the water cycle above the north polar cap on Mars using MRO CRISM retrievals of water vapor. Icarus 321, 722–735 (2019). https:// doi.org/10.1016/j.icarus.2018.12.024 Khayat, A.S., Smith, M.D., Wolff, M. et al.: ExoMars TGO/NOMAD-UVIS vertical profiles of ozone: 2 the high-altitude layers of atmospheric ozone. J. Geophys. Res. Planets 126(11), e2021JE006834 (2021) Korablev, O., et al.: No detection of methane on Mars from early ExoMars Trace Gas Orbiter observations. Nat. 568(7753), 517–520 (2019) Krasnopolsky, V.A.: Search for methane and upper limits to ethane and SO2 on Mars. Icarus 217(1), 144–152 (2012) Krasnopolsky, V.A., Korablev, O.I., Moroz, V.I., et al.: Infrared solar occultation sounding of the Martian atmosphere by the Phobos spacecraft. Icarus 94(1), 32–44 (1991) Lebonnois, S., Quémerais, E., Montmessin, F., et al.: Vertical distribution of ozone on Mars as measured by SPICAM/Mars express using stellar occultations. J. Geophys. Res. Planets 111(E9) (2006) Maltagliati, L., Montmessin, F., Korablev, O., et al.: Annual survey of water vapor vertical distribution and water–aerosol coupling in the martian atmosphere observed by SPICAM/MEx solar occultations. Icarus 223(2), 942–962 (2013) Montmessin, F., Lefèvre, F.: Transport-driven formation of a polar ozone layer on Mars. Nat. Geosci. 6(11), 930–933 (2013) Nier, A.O., McElroy, M.B.: Composition and structure of Mars’ upper atmosphere: results from the neutral mass spectrometers on Viking 1 and 2. J. Geophys. Res. 82(28), 4341–4349 (1977) Ojha, L., Karunatillake, S., Iacovino, K.: Atmospheric injection of sulfur from the Medusae Fossae forming events. Planet Space Sci. 179, 104734 (2019) Patel, M.R., Sellers, G., Mason, J.P., et al. ExoMars TGO/NOMAD-UVIS vertical profiles of ozone: 1 seasonal variation and comparison to water. J. Geophys. Res. Planets 126(11), e2021JE006837 (2021) Perrier, S., Bertaux, J.L., Lefèvre, F. et al.: Global distribution of total ozone on Mars from SPICAM/ MEX UV measurements. J. Geophys. Res. Planets 111(E9) (2006) Pla-Garcia, J., Rafkin, S.C., Karatekin, Ö., Gloesener, E.: Comparing MSL curiosity rover TLSSAM methane measurements with Mars regional atmospheric modeling system atmospheric transport experiments. J. Geophys. Res. Planets 124(8), 2141–2167 (2019) Rodin, A.V., Korablev, O.I., Moroz, V.I.: Vertical distribution of water in the near-equatorial troposphere of Mars: water vapor and clouds. Icarus 125(1), 212–229 (1997) Stern, J.C., Sutter, B., Freissinet, C., et al.: Evidence for indigenous nitrogen in sedimentary and aeolian deposits from the Curiosity rover investigations at Gale crater. Mars Proc. Natl. Acad. Sci. 112(14), 4245–4250 (2015) Stevens, M.H., et al.: Detection of the nitric oxide dayglow on Mars by MAVEN/IUVS. J. Geophys. Res. Planet 124(5): 1226–1237 (2019) Vandaele, A.C., Daerden, F., Thomas, I., et al.: Impact of the 2018 global dust storm on Mars atmosphere composition as observed by NOMAD on ExoMars Trace Gas Orbiter. In: AGU Fall Meeting Abstracts, vol. 2018, pp. P31A-03 (2018) Webster, C.R., Mahaffy, P.R., Atreya, S.K., et al.: Mars methane detection and variability at gale crater. Science 347(6220), 415–417 (2015) Webster, C.R., Mahaffy, P.R., Atreya, S.K., et al.: Background levels of methane in Mars’ atmosphere show strong seasonal variations. Science 360(6393), 1093–1096 (2018)

Chapter 20

Seasonal Variability of Atmospheric Gases

Abstract Since the axial tilt 25.2° of Mars is similar to the Earth’s tilt 23.5°, the atmosphere of Mars also experiences strong seasonal variations. The seasonal cycle of Mars is dominated by a strong and asymmetric Hadley cell, extending between north and south mid-latitudes producing trade winds and allowing cross-equatorial transport of dust and trace gases. The neutral densities of various atmospheric gases (O3 , H2 O, CO2 , O2 and CO) are estimated in the Martian atmosphere by global model GCM. Keywords Seasons variability · Dust · Lower atmospheric gases

20.1 Basics of Global Circulation Model (GCM) A GCM for Mars has been developed by LMD (Laboratoire de Météorologie Dynamique), France (Hourdin et al. 1993) which simulates atmospheric compositions. The GCM is developed by using the time-dependent, coupled Navier–Stokes, continuity, momentum and Boltzmann equations. Using Eq. (10.1) the Boltzmann transport equation can be written as follows: ∂ F→S + v→ · ∇r F→S + a · ∇v F→S − Q(→ r , v→, t) − ∂t

(

δ F→S δt

) =0

(20.1)

Coll.

where ∂ F→s /∂t represents local variation of distribution function F→s of species s with time t, ∇r is the space derivative, ∇v is the velocity derivative, ν→ is the velocity, and a is the acceleration. Therefore, the second term on the left-hand side describes the change in F→s due to spatial gradients, and the third term describes the presence of external forces, Q)is the source term for production rates obtained from solar radia( represents chemical production/loss rates. In order to deduce tion, and ∂ F→s /∂t Coll the transport equations of macroscopic parameters one first defines velocity moments of this distribution function. The most commonly used macroscopic parameters of interest are: © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_20

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∫∞ n S (r, t) =

d 3 υS f S

(20.2)

d 3 υ S f S ν S /n S

(20.3)

mS 3 d υ S f S (ν S − u→ S )2 3kn S

(20.4)

−∞

∫∞ u→ S (r, t) = −∞

TS = mS q→S = 2

∫∞ d 3 υ S f S (ν S − u→s )2 (ν S − u→s )

(20.5)

−∞

mS pS = 3

∫∞ d 3 υ S f S (ν S − u→ S )2

(20.6)

d 3 υ S f S (ν S − u→ S )(ν S − u→ S )

(20.7)

−∞

∫∞

P→S = m S

−∞

τ S = P→S − p S /I

(20.8)

where ns is the number density of species s, us is the drift (mean) velocity, Ts is the temperature, qs is the heat flow vector, ps is the scalar pressure, Ps is the pressure tensor and τs is the stress tensor.

20.2 Ozone Variability Global distribution of ozone (longitudinal averaged) was reported from remote sensing observations by Perrier et al. (2006) and Clancy et al. (2016). The SPICAM on board MEX and NOMAD on board TGO have observed seasonal variability of ozone density in the atmosphere of Mars at different latitudes and longitudes (Montmessin et al. 2017; Patel et al. 2021). Recently in Fig. 20.1, Haider et al. (2022) examined the seasonal and latitudinal variability of ozone during the dust storm and non-dust storm periods. They reported that during the thick layer of dust, the sunlight cannot penetrate deep into the atmosphere of Mars. This interrupts the dissociation of ozone below the dust layer, which is the main loss mechanism of ozone in dayside atmosphere. On the other hand, production of O3 is directly related to the production of the oxygen atom, which is produced more above the dust layer from the photolysis of CO2 . Therefore, the O3 enhanced during the dust storm period.

20.3 Global Mapping of Water Vapour

181

Fig. 20.1 The seasonal variability of zonally averaged column ozone at a 10–30 N, b 30–50 N, c 50–70 N, d 10–30 S, e 30–50 S, f 50–70 S observed by SPICAM and modelled by MCD between MY27 and MY30. The effect of dust storms on column. Ozone at latitude 10–30 S is shown by an arrow (from Haider et al. 2022)

20.3 Global Mapping of Water Vapour Water vapour is one of the most variable trace species and is important for the stability of the Martian atmosphere. It is the only source of the catalytic species HOx in Martian atmosphere. The first measurement of water vapour was carried out by MAWD on board Viking orbiters (Farmer et al. 1977; Jakosky and Farmer 1982). Later systematic mappings of water vapour were carried out from TES on board MGS (Smith 2002, 2004; Pankine et al. 2010), the PFS, OMEGA and SPICAM instruments on board MEX (Fedorova et al. 2006; Melchiorri et al. 2007; Maltagliati et al. 2011, 2013), CRISM on board MRO (Smith et al. 2009) and NOMAD on board TGO (Vandaele et al. 2019).

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Figure 20.2 shows the seasonal variation of the water vapour vertical profiles retrieved from NOMAD instrument (Aoki et al. 2022). This observation shows a strong contrast between aphelion and perihelion periods. In the northern hemisphere (top Fig. 22.3) the significant increase of the water vapour reached up to 100 km at Ls ~ 190° in MY34. In the southern hemisphere (bottom Fig. 22.3) an increase of water vapor abundances due to dust storm is also observed up to 80 km at Ls ~ 270°. The previous studies of GCM confirmed that water vapour can be transported to such high altitudes (Daerden et al. 2019; Holmes et al. 2021; Shaposhnikov et al. 2019).

20.4 Seasonal Variability of CO2 Carbon dioxide is the major component of the Martian atmosphere, with large seasonal variations in the global surface density due to the condensation and sublimation of CO2 in the polar region during winter and spring, respectively (Tillman et al. 1993). This measurement was first carried out by Viking Landers through global fluctuations of about 30% in surface pressure. After about 3–4 decades SAM instrument on board MSL Curiosity rover observed the CO2 mixing ratios and pressure between MY31 and MY34 (Trainer et al. 2019). Figure 20.3 represents the SAM measurements of mixing ratios and pressure cycles of CO2 on the surface of Mars. In this figure CO2 is modulated with a lag of approximately 20°–40° of Ls behind the maxima and minima of the total pressure cycle.

20.5 Seasonal Variability of N2 , Ar, O2 and CO The mixing ratios of Ar, N2 and O2 observed by SAM instrument are shown in Fig. 20.4. This instrument has also observed the mixing ratio of CO, which is plotted in Fig. 20.5. These four gases are not condensable at Mars surface and atmospheric temperatures and pressures and thus are not expected to deposit or sublime from the polar caps as does CO2 . However, there are seasonal trends in the mixing ratios of these gases that are a response to the mixing of air masses during the seasonal CO2 cycle. The CO is also observed from ground based observations at infrared, millimeter and submillimeter wavelengths (e.g. Clancy et al. 1983, 1990; Lellouch et al. 1991; Krasnopolsky 2003, 2007; Moreno et al. 2009). The CO mixing ratio is also observed for certain locations and seasons from the OMEGA (Encrenaz et al. 2006) and PFS (Billebaud et al. 2009; Sindoni et al. 2011) instruments on board MEX. The complete climatology of CO mixing ratio was observed using near infrared observations made by the CRISM instrument on board the MRO (Smith et al. 2018). The ground based and satellite observations of CO mixing ratios were greater than the insitu values (Sindoni et al. 2011).

Fig. 20.2 Seasonal variability of ozone in northern hemisphere (top) and southern hemisphere (bottom) observed by NOMAD (from Aoki et al. 2022)

20.5 Seasonal Variability of N2 , Ar, O2 and CO 183

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Fig. 20.3 SAM measurements of CO2 volume mixing ratios and pressure during MY31-MY34 at Ls 0–360o (from Trainer et al. 2019)

Fig. 20.4 Seasonal variability of the mixing ratios of N2 , Ar and O2 observed by SAM instrument during MY31-MY34 (from Trainer et al. 2019)

References

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Fig. 20.5 Seasonal variability of CO mixing ratio observed by SAM instrument during MY31 and MY32 (from Trainer et al. 2019)

References Aoki, S., Gkouvelis, L., Gérard, J.C., et al.: Density and temperature of the upper mesosphere and lower thermosphere of Mars retrieved from the OI 5577 nm dayglow measured by TGO/ NOMAD. J. Geophys. Res. Planets e2022JE007206 (2022) Billebaud, F., Brillet, J., Lellouch, E., et al.: Observations of CO in the atmosphere of Mars with PFS onboard Mars express. Planet Space Sci. 57(12), 1446–1457 (2009) Clancy, R.T., Muhleman, D.O., Jakosky, B.M.: Variability of carbon monoxide in the Mars atmosphere. Icarus 55(2), 282–301 (1983) Clancy, R.T., Muhleman, D.O., Berge, G.L.: Global changes in the 0–70 km thermal structure of the Mars atmosphere derived from 1975 to 1989 microwave CO spectra. J. Geophys. Res. Solid Earth 95(B9), 14543–14554 (1990) Clancy, R.T., Wolff, M.J., Lefèvre, F. et al.: Daily global mapping of Mars ozone column abundances with MARCI UV band imaging. Icarus 266, 112–133 (2016) Daerden, F., Neary, L., Viscardy, S., et al.: Mars atmospheric chemistry simulations with the GEMMars general circulation model. Icarus 326, 197–224 (2019) Encrenaz, T., Fouchet, T., Melchiorri, R., et al.: Seasonal variations of the martian CO over Hellas as observed by OMEGA/Mars express. Astron. Astrophys. 459(1), 265–270 (2006) Farmer, C.B., Davies, D.W., Holland, A.L., et al.: Mars: Water vapor observations from the Viking orbiters. J. Geophys. Res. Space 82(28), 4225–4248 (1977) Fedorova, A., Korablev, O., Bertaux, J.L.: Mars water vapor abundance from SPICAM IR spectrometer: seasonal and geographic distributions. J. Geophys. Res. Planets 111(E9) (2006) Haider, S.A., Shah, S.Y., Masoom, J., Sheel, V., Kuroda, T.: Impact of dust loading on ozone, winds and heating rates in the atmosphere of mars: seasonal variability, climatology and SPICAM observations. Planet Space Sci. 212, 105424 (2022) Holmes, J.A., Lewis, S.R., Patel, M.R., et al.: Enhanced water loss from the martian atmosphere during a regional-scale dust storm and implications for long-term water loss. Earth Planet. Sci. Lett. 571, 117109 (2021) Holsclaw, G.M., Deighan, J., Almatroushi, H., et al.: The Emirates Mars Ultraviolet Spectrometer (EMUS) for the EMM mission. Space Sci. Rev. 217(8), 1–49 (2021) Hourdin, F., Van, P.L., Forget, F., Talagrand, O.: Meteorological variability and the annual surface pressure cycle on Mars. J. Atmos. Sci. 50(21), 11 (1993)

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Jakosky, B.M., Farmer, C.B.: The seasonal and global behavior of water vapor in the Mars atmosphere: complete global results of the Viking atmospheric water detector experiment. J. Geophys. Res. Solid Earth 87(B4), 2999–3019 (1982) Krasnopolsky, V.A.: Long-term spectroscopic observations of Mars using IRTF/CSHELL: mapping of O2 dayglow, CO, and search for CH4 . Icarus 190(1), 93–102 (2007) Krasnopolsky, V.A.: Spectroscopic mapping of Mars CO mixing ratio: detection of north-south asymmetry. J. Geophys. Res. Planets 108(E2) (2003) Lellouch, E., Paubert, G., Encrenaz, T.: Mapping of CO millimeter-wave lines in Mars’ atmosphere: the spatial variability of carbon monoxide on Mars. Planet Space Sci. 39(1–2), 219–224 (1991) Maltagliati, L., Titov, D.V., Encrenaz, T., et al.: Annual survey of water vapor behavior from the OMEGA mapping spectrometer onboard Mars express. Icarus 213(2), 480–495 (2011) Maltagliati, L., Montmessin, F., Korablev, O., et al.: Annual survey of water vapor vertical distribution and water–aerosol coupling in the martian atmosphere observed by SPICAM/MEx solar occultations. Icarus 223(2), 942–962 (2013) Maltagliati, L., Montmessin, F., Korablev, O. et al.: Annual survey of water vapor vertical distribution and water–aerosol coupling in the martian atmosphere observed by SPICAM/MEx solar occultations. Icarus 223(2), 942–962 (2013) Melchiorri, R., Encrenaz, T., Fouchet, T., et al.: Water vapor mapping on Mars using OMEGA/Mars express. Planet Space Sci. 55(3), 333–342 (2007) Montmessin, F., Korablev, O., Lefèvre, F., Bertaux, J.L., Fedorova, A., Trokhimovskiy, A., Chapron, N.: SPICAM on mars express: a 10 year in-depth survey of the Martian atmosphere. Icarus 297, 195–216 (2017) Moreno, R., Lellouch, E., Forget, F., et al.: Wind measurements in Mars’ middle atmosphere: IRAM Plateau de Bure interferometric CO observations. Icarus 201(2), 549–563 (2009) Pankine, A.A., Tamppari, L.K., Smith, M.D.: MGS TES observations of the water vapor above the seasonal and perennial ice caps during northern spring and summer. Icarus 210(1), 58–71 (2010) Patel, M.R., Sellers, G., Mason, J.P., et al.: ExoMars TGO/NOMAD-UVIS vertical profiles of ozone: 1 seasonal variation and comparison to water. J. Geophys. Res. Planets 126(11), e2021JE006837 (2021) Perrier, S., Bertaux, J.L., Lefèvre, F., et al.: Global distribution of total ozone on Mars from SPICAM/ MEX UV measurements. J. Geophys. Res. Planets 111(E9) (2006) Shaposhnikov, D.S., Medvedev, A.S., Rodin, A.V., Hartogh, P.: Seasonal water “pump” in the atmosphere of Mars: vertical transport to the thermosphere. Geophys. Res. Lett. 46(8), 4161– 4169 (2019) Sindoni, G., Formisano, V., Geminale, A.: Observations of water vapour and carbon monoxide in the Martian atmosphere with the SWC of PFS/MEX Planet. Space Sci. 59(2–3), 149–162 (2011) Smith, M.D.: The annual cycle of water vapor on Mars as observed by the thermal emission spectrometer. J. Geophys. Res. Planets 107(E11), 25–31 (2002) Smith, M.D.: Interannual variability in TES atmospheric observations of Mars during 1999–2003. Icarus 167(1), 148–165 (2004) Smith, M.D.: THEMIS observations of Mars aerosol optical depth from 2002–2008. Icarus 202(2), 444–452 (2009) Smith, M.D., Daerden, F., Neary, L., Khayat, A.: The climatology of carbon monoxide and water vapor on Mars as observed by CRISM and modeled by the GEM-Mars general circulation model. Icarus 301, 117–131 (2018) Tillman, J.E., Johnson, N.C., Guttorp, P., Percival, D.B.: The Martian annual atmospheric pressure cycle: years without great dust storms. J. Geophys. Res. Planets 98(E6), 10963–10971 (1993) Trainer, M.G., Wong, M.H., McConnochie, T.H., et al.: Seasonal variations in atmospheric composition as measured in Gale Crater. Mars. J. Geophys. Res. Planets 124(11), 3000–3024 (2019) Vandaele, A.C., Korablev, O., Daerden, F., et al.: Martian dust storm impact on atmospheric H2 O and D/H observed by ExoMars trace gas orbiter. Nature 568(7753), 521–525 (2019)

Chapter 21

Infrared Thermal Emissions from Mars Atmosphere

Abstract The infrared spectroscopy is an important remote sensing technique to study the planetary atmospheres and surfaces. Over several decades this technique has been used on Mars to study the composition of the surface, minerals, dust, water ice clouds, temperature and atmospheric gases (Hanel et al., Icarus 17:423–442, 1972; Kieffer et al., J. Geophys. Res. 82:4249–4291, 1977; Christensen et al., J. Geophys. Res. Planets 106:23,873–23,885, 2001). The Martian atmosphere consist CO2 bands, water vapour lines, water ice clouds, and signatures of dust periodically stirred up by strong surface winds. The surface of Mars has been observed in certain spectral windows: 780–1000 cm−1 , 1080–1240 cm−1 and 2500–2800 cm−1 .The presence of water and dust on Martian surface have been observed by Mariner 9, Viking, MGS, Mars Odyssey and MEX from infrared thermal emission spectrometer. (Hanel et al., Icarus 17:423–442, 1972; Kieffer et al., J. Geophys. Res. 82:4249–4291, 1977; Christensen et al., J. Geophys. Res. Planets 106:23,873–23,885, 2001). These missions have also observed surface minerals, rocks and temperatures on Mars. Keywords PFS observations · Thermal emissions · Brightness

21.1 PFS Observations The PFS instrument onboard MEX observed thermal emission spectra emitted by the surface and atmosphere of Mars. This instrument has a short-wavelength (SW) channel covering the spectral range 1700–8200 cm−1 (1.2–5.5 μm) and long wavelength (LW) channel covering 250–1700 cm−1 (5.5–45 μm) (Giuranna et al. 2005a, b; Fouchet et al., 2008; Wolkenberg et al. 2011). Formisano et al. (2005) was first to discover methane in SW channel of PFS spectra at wave number 3018 cm−1 . The global average methane mixing ratio ~ 10 ± 5 ppbv have been reported by them. Later seasonal, diurnal and spatial variations of methane were studied by Geminale et al. (2008, 2011) in the atmosphere of Mars. Recently the mixing ratios of water vapour and carbon monoxide have been observed by Sindoni et al. (2011) at wave numbers 3845 cm−1 (2.6 μm) and 4235 cm−1 (2.36 μm) respectively. They reported that there is anti-correlation between the water vapour and CO concentrations. This is because © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_21

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of sublimation of the CO2 and water vapour from north polar ice cap, which decreases concentration of non-condensable species CO. In the south polar cap similar process also occurred but in this case the condensation of CO2 and water vapour increases the concentration of non-condensable species CO. Fouchet et al. (2008) used PFS data of LW channel to study the seasonal and geographical variations of water vapour for MY26 and MY27. They observed two maxima in the geographical distributions of water vapour over Tharsis and Arabia regions. The mechanisms of CO2 condensation and accumulation over south polar cap were studied by Giuranna et al. (2008). They reported that polar cap expands symmetrically during the fall season with a constant speed. Zasova et al. (2006) suggested that the atmospheric dust has a significant impact on radiative heating. Later Mättänen et al. (2009) analysed PFS data to study the response of local dust storm on atmospheric thermal structure near the equator of Mars in the late northern summer. They reported that during the dust storm period the temperature near the surface were ~ 10 K colder and ~ 5 K warmer in the upper atmosphere. Sato et al. (2011) investigated tidal variations in the atmospheric temperature at low altitude (< 45 km) during the dust-clear period from PFS data. Using MGS and Odyssey data, the inter-annual and seasonal variabilitiy of dust optical depths were studied by Montabone et al. (2015) between MY24 and MY32. analysed the data obtained from MGS and Odyssey for MY24 to MY32. Aoki et al. (2015) have reported H2 O2 in the Martian atmosphere at wave number 379 cm−1 . They have derived mixing ratios 16 ± 19 ppb at Ls = 0°–120°, 35 ± 32 ppb at Ls = 120°–240° and 41 ± 28 ppb at Ls = 240°–360°.

21.2 PFS Data Analysis Masoom et al. (2019) have analysed PFS data of LW channel for MY28 when a major dust storm occurred at low latitudes in southern hemisphere between Ls = 260°–300° for about a couple of weeks. They retrieved thermal emission spectra and brightness temperature from these observations in presence and absence of dust storm. They have considered that the thermal emissions emitted from the surface of Mars in the infrared wavelength v, under local thermal equilibrium conditions is characterized by Planck function (Haus and Titov 2000) B and it is expressed as: 2hc2 v3 } B(T) = { hcv e KT − 1

(21.1)

where v is the wave number and T is temperature, h is Planck constant (6.62 × 10–34 J.S), k is Boltzmann constant (1.38 × 10–34 J.K−1 ), and c is speed of light (3.00 × 108 m.s−1 ). The altitude dependent brightness temperatures (Haus and Titov 2000) are calculated by inverting above equation:

21.3 Brightness Temperature of Mars

189

TB = k ln

hcv ( 2 3) 2hc v BT +1

(21.2)

The radiative transfer equation describes that the radiation is affected through the Martian atmosphere. The emission, absorption and scattering processes are considered in the radiative transfer equation (Ignatiev et al. 2005). Thermal emission is characterized by the Planck function B. In Sects. 23.4 and 23.5 Planck function and radiative transfer model are described in detail. The modeled results are presented in this chapter. The description of the PFS data processing is shown in the Fig. 21.1 The IDL routines provided by the ESA, has been used to process the raw data files. The routine “LWC_CAL.PRO” and Geometry_READER.PRO reads raw data of PFS instrument and the spacecraft geometry respectively. We select LWC_CAL.PRO routine. This routine processes the raw data along with two HK files and corresponding six LABEL files. Later we select the path DATA directory Geometry_ READER.PRO and give orbit number ‘XXXX’. This routine is computed, whose output is obtained as radiance spectra versus wave numbers at Spacecraft Elapsed Time (SCET) for each selected ORBIT. The orbit’s geometry of MEX spacecraft can be known by processing GEO files of the orbit. This routine provides the SpaceCraft Elapsed Time (SCET) and geometrical information (MY, Ls, Latitude, and Longitude etc.). Like other instruments, PFS instrument also observed instrumental error while measuring thermal emission spectra. This error depends upon the measurement conditions and mechanical and non-mechanical vibration of the spacecraft (Formisano et al. 2005). Comolli and Saggin (2010) showed that the uncertainties can be reduced by the averaging of the spectra. Giuranna et al. (2005a; b) reported about 1–σ error less than 1% in PFS measurements. These PFS data are averaged over 10 degree latitudes between 0° to 30°S to decrease the instrumental noise (Comolli and Saggin 2010). The radiative transfer modeling of these measurements are carried out at Ls = 240°, 280°, 300° and 320° due to absorptions of CO2 , H2 O and dusts in southern hemisphere at low latitudes (Masoom et al. 2019).

21.3 Brightness Temperature of Mars The Fig. 21.2a–c represents the thermal emission measurements carried out by PFS between wave numbers 250–1400 cm−1 at latitude range 0°–10o S, 10°–20o S and 20°–30° S respectively. These spectra were observed in MY28 at Ls = 240°, 280°, 300° and 320° corresponding to orbit # 4338, 4552, 4670 and 4808 respectively (The emission spectra is not observed at Ls = 320o between latitude range 0–10° S). In these spectra prominent dip at about 667 cm−1 is observed due to absorption of CO2 . The Fig. 21.3a–c represents the brightness temperature profile between wave number 250–1400 cm−1 at latitude 0°–10° S, 10°–20° S and 20°-30° S respectively. The brightness temperatures are obtained from the thermal emission spectra (shown

190

21 Infrared Thermal Emissions from Mars Atmosphere

Fig. 21.1 Schematic diagram for processing of raw data of PFS measurements and spacecraft geometry using IDL package and other data/geometry routines (from Masoom et al. 2019)

Radiance (erg/(cm2 sec ster cm-1)) Radiance (erg/(cm 2 sec ster cm-1))

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21.3 Brightness Temperature of Mars

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21 Infrared Thermal Emissions from Mars Atmosphere

in Fig. 23.2a–c) by inverting Planck function. In the invert calculation, non-linearity of Planck function will be eliminated. Therefore, the brightness temperature is nearly same at wave number ≤ 600 cm−1 . There is decrease in brightness temperature due to absorption of dust between wave number 900–1200 cm−1 . When the dust is raised into the atmosphere the daytime temperature decreases because atmospheric layers are cooler than the surface of Mars (Montabone et al. 2015; Wolkenberg et al. 2011). The major dip in the brightness temperature due to absorption of CO2 is observed at wave number 667 cm−1 . The temperature is decreasing with increasing Ls between wave number 250– 1400 cm−1 . The maximum temperature 260–280 K is observed at Ls = 240° when Mars reaches at perihelion and it receives a large amount of solar radiation. The dust storm is observed between summer onset and mid-summer at Ls = 280° and 300° respectively when Mars was moving away from perihelion. The temperature decreases to ~ 220 K during the late summer at Ls = 320°. In this season a small dip occurred in the temperature at wave number 1300 cm−1 , which is produced due to low emissivity in presence of larger particles of mineral dust near Christiansen frequency (Formissano et al. 2005).

21.4 Planck Function and Radiance The observed thermal emission spectra are characterized by Planck function at the corresponding surface temperature under Local Thermodynamic Equilibrium condition (LTE) (Haus and Titov 2000). The Planck function is expressed as given below: BT = {

2hc2 v 3

} hcv e KT − 1

(21.3)

where BT is the observed thermal emission spectra at brightness temperature T, υ is the wave number, h is Planck constant, k is Boltzmann constant, and c is speed of light. Sometimes the Eq. (21.3) is expressed into two constants, c1 and c2 . These constants are called as radiation constants. The first radiation constant c1 = 2π hc2 = 3.74 × 1016 W m2 . The second radiation constant c2 = hc/k = 1.44 × 102 m K. The Planck function is varying with altitudes because the temperature T is altitude dependent. The altitude-dependent of brightness temperature (Haus and Titov 2000) can be calculated by inverting Eq. (21.3) as follows: TB = k ln

hcv ( 2 3) 2hc v BT +1

(21.4)

In the above Eqs. (21.3) and (21.4) we have now considered molecular scattering under LTE condition. Therefore these equations cannot calculate the spectral features

21.4 Planck Function and Radiance

193

Brightness Temperature (K)

320 280 240 200 Orbit 4338, Ls = 240o Orbit 4552, Ls = 280o

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(c) 120 250

450

650

850

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1250

1450

Fig. 21.3 The brightness temperature inverted from radiance observed by PFS instrument at latitudes: a 0°–10° S, b 10°–20° S and c 20°–30° S (from Masoom et al. 2019)

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21 Infrared Thermal Emissions from Mars Atmosphere

of thermal emission spectrum. The molecular scattering is considered in the radiative transfer method.

21.5 Radiative Transfer Model The radiative transfer equation describes that the radiation is affected through the emission, absorption and scattering processes (Ignatiev et al. 2005; Masoom et al. 2019). These three processe are coupled and can be represented as: d Iλ = −(aλ (s) + dλ (s))Iλ (s) + aλ (s)Bλ (T ) ds ∫ ) ( dλ + P Ω, Ω' Iλ (Ω' )dΩ' 4π

(21.5)

The sum aλ + dλ explain the extinction coefficient and can be expressed in terms of atmospheric optical depth in vertical direction s: dτλ = (aλ (s) + dλ (s))ds

(21.6)

Thus dIλ (τ) wd = −Iλ (τ) + wa Bλ (T) + dτ 4π

∫

) ( p Ω, Ω' Iλ (Ω' )dΩ'

(21.7)

where, wa = aλ /(aλ + dλ ) and wd = dλ /(aλ + dλ ) are absorption and scattering albedo respectively. By using above equations we get Iλ (τ ) = Iλ (0)e

−τ

∫τ +

'

Bλ (T )e−(τ −τ ) dτ '

(21.8)

0

In Eq. (21.8) first term is a pure extinction term while second term describes the emission from the atmosphere. Rayleigh scattering due to molecules is significantly low in the Martian atmosphere because of its λ−4 dependence (Haus and Titov 2000). Therefore, Rayleigh scattering is not included in Eq. (21.8). Iλ (0) is the emitted radiation from the surface of Mars at wavelength λ. The atmospheric contributions have been incorporated at every height. The contributions emitted from the elements of distance dτ at different τ’ along the path are themselves attenuated by the exponential factor over (τ τ’). The optical depth τ is defined as: ∫s τ (s) =

∫s Nco2 σco2 (λ)ds +

0

NH2 o σH2 o (λ)ds 0

21.6 Modeling of Thermal Emission Spectra

195

∫s +

Ndust σdust (λ)ds

(21.9)

0

where σ represents the absorption cross sections of CO2 , H2 O and dust in the infrared wavelength range 250–1400 cm−1 . The cross sections for CO2 and H2 O are taken from Exomol Spectroscopic Database (Tennyson and Yurchenko 2012). The cross section of dust is taken from Conrath et al. (1973). N represents the number density of CO2 , H2 O and dust. The number densities of CO2 and H2 O are taken from MCD v5.2 (Millour et al. 2014) for Ls = 240°, 280°, 300o and 320° at low latitude region averaged over 0°–10° S, 10°–20° S and 20°–30° S The dust density profiles are calculated using the profile shape parameters obtained from recent observations carried out by MCS onboard MRO (Heavens et al. 2014; Guzewich et al. 2014).

21.6 Modeling of Thermal Emission Spectra Using Eqs. 21.5–21.9 thermal emission spectra were estimated between wavelength range 200–1400 cm−1 at low latitude range 0°–10o S, 10°–20° S, and 20°–30° S for Ls = 240°, 280°, 300° and 320°, when a major dust storm occurred in southern hemisphere during MY28. The cross section and neutral density were taken from Conrath et al. (1973) and Millour et al. (2014) respectively. In Fig. 21.4a–c the measured and estimated thermal emission spectra are compared. The model reproduces the absorption features of CO2 at wave number 600–750 cm−1 .

21 Infrared Thermal Emissions from Mars Atmosphere

Radiance (erg/(cm2 sec ster cm-1))

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Wavenumber (cm-1) Fig. 21.4 Comparison of estimated thermal emission spectra with the measurements carried out by PFS onboard MEX at latitudes: a 0°–10° S, b 10°–20° S, and c 20°–30° S (from Masoom et al. 2019)

References

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References Aoki, S., Nakagawa, H., Sagawa, H., et al.: Seasonal variation of the HDO/H2 O ratio in the atmosphere of Mars at the middle of northern spring and beginning of northern summer. Icarus 260, 7–22 (2015) Christensen, P.R., Morris, R.V., Lane, M.D., et al.: Global mapping of Martian hematite mineral deposits: remnants of water-driven processes on early Mars. J. Geophys. Res. Planets 106(E10), 23873–23885 (2001) Comolli, L., Saggin, B.: Analysis of disturbances in the Planetary Fourier Spectrometer through numerical modeling. Planet Space Sci. 58(5), 864–874 (2010) Conrath, B., Curran, R., Hanel, R., et al.: Atmospheric and surface properties of Mars obtained by infrared spectroscopy on Mariner 9. J. Geophys. Res. 78(20), 4267–4278 (1973) Formisano, V., Angrilli, F., Arnold, G., et al.: The planetary Fourier spectrometer (PFS) onboard the European Mars Express mission. Planet Space Sci. 53(10), 963–974 (2005) Fouchet, T., Lellouch, E., Ignatiev, N., et al.: Martian water vapour: Mars Express PFS/LW observations and comparison with LMD/GCM simulations. 37th COSPAR Sci. Assembly 37, 921 (2008) Geminale, A., Formisano, V., Giuranna, M.: Methane in Martian atmosphere: average spatial, diurnal, and seasonal behavior. Planet Space Sci. 56(9), 1194–1203 (2008) Geminale, A., Formisano, V., Sindoni, G.: Mapping methane in Martian atmosphere with PFS-MEX data. Planet Space Sci. 59(2–3), 137–148 (2011) Giuranna, M., Grassi, D., Formisano, V., et al.: PFS/MEX observations of the condensing CO2 south polar cap of Mars. Icarus 197(2), 386–402 (2008) Giuranna, M., Formisano, V., Biondi, D., et al.: Calibration of the Planetary Fourier Spectrometer short wavelength channel. Planet Space Sci. 53(10), 975–991 (2005a) Giuranna, M., Formisano, V., Biondi, D. et al.: Calibration of the Planetary Fourier Spectrometer long wavelength channel. Planet Space Sci. 53(10), 993–1007 (2005b) Guzewich, S.D., Smit, M.D., Wolff, M.J.: The vertical distribution of Martian aerosol particle size. J. Geophys. Res. Planets 119(12), 2694–2708 (2014) Hanel, R., Conrath, B., Hovis, W., et al.: Investigation of the Martian environment by infrared spectroscopy on Mariner 9. Icarus 17(2), 423–442 (1972) Haus, R., Titov, D.V.: PFS on Mars Express: preparing the analysis of infrared spectra to be measured by the Planetary Fourier Spectrometer. Planet Space Sci. 48(12–14), 1357–1376 (2000) Heavens, N.G., Johnson, M.S., Abdou, W.A., et al.: Seasonal and diurnal variability of detached dust layers in the tropical Martian atmosphere. J. Geophys. Res. Planets 119(8), 1748–1774 (2014) Ignatiev, N.I., Grassi, D., Zasova, L.V.: Planetary Fourier spectrometer data analysis: fast radiative transfer models. Planet Space Sci. 53, 1035 (2005). https://doi.org/10.1016/jpss200412009 Kieffer, H.H., Martin, T.Z., Peterfreund, A.R., et al.: Thermal and albedo mapping of Mars during the Viking primary mission. J. Geophys. Res. 82(28), 4249–4291 (1977) Määttänen, A., Fouchet, T., Forni, O., et al.: A study of the properties of a local dust storm with Mars Express OMEGA and PFS data. Icarus 201(2), 504–516 (2009) Masoom, J., Haider, S.A., Giuranna, M.: Response of dust on thermal emission spectra observed by Planetary Fourier Spectrometer (PFS) on-board Mars Express (MEX). Indian J. of Radio Space Phys, vol. 48, 38-44 (2019) Millour, E., Forget, F., Spiga, A., et al.: A new Mars climate database v5 1. In: The fifth international workshop on the Mars atmosphere: modelling and observations, pp. id-1301 (2014) Montabone, L., Forget, F., Millour, E., Wilson, R.J., Lewis, S.R., Cantor, B., Wolff, M.J.: Eight-year climatology of dust optical depth on Mars. Icarus 251, 65–95 (2015) Sato, T.M., Fujiwara, H., Takahashi, Y.O., et al.: Tidal variations in the Martian lower atmosphere inferred from Mars Express Planetary Fourier Spectrometer temperature data. Geophys. Res. Lett. 38(24) (2011)

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Sindoni, G., Formisano, V., Geminale, A.: Observations of water vapour and carbon monoxide in the Martian atmosphere with the SWC of PFS/MEX Planet. Space Sci. 59(2–3), 149–162 (2011) Tennyson, J., Yurchenko, S.N.: ExoMol: molecular line lists for exoplanet and other atmospheres. Mon. Not. R. Astron. Soc. 425(1), 21–33 (2012) Wolkenberg, P., Smith, M.D., Formisano, V., Sindoni, G.: Comparison of PFS and TES observations of temperature and water vapor in the martian atmosphere. Icarus 215(2), 628–638 (2011) Zasova, L.V., Formisano, V., Moroz, V.I., et al.: Results of measurements with the Planetary Fourier Spectrometer onboard Mars Express: clouds and dust at the end of southern summer A comparison with OMEGA images. Cosmic Res. 44(4), 305–316 (2006)

Chapter 22

Lower Ionosphere of Mars

Abstract The lower ionosphere is produced due to following sources: (1) solar Lyman α (1216 Å) ionizing the minor constituent NO, (2) solar X-ray (λ < 8 Å) ionizing N2 and O2 , (3) cosmic ray ionizing all atmospheric gases and (4) photoionization of the metastable O2 (1 Δg) by solar UV radiation (λ < 1118 Å). In the lower ionosphere of Mars the D region is produced at about 30 km due to impact of GCR and hard X-rays (0.5–3 Å). GCR flux is exponentially attenuated in the lower atmosphere and is calculated between values 103 –10−5 particles m−2 s−1 GeV−1 ster−1 at energy range of 1–1000 GeV (Molina-Cuberos et al., J. Geophys. Res. Planets 107:3–1, 2002; Haider et al., J. Geophys. Res. Space Phys. 114, 2009). The flux of GCR depends on the solar activity which decelerates it in the interplanetary medium. This flux has been used by several investigators for average condition to study the lower ionosphere of Mars (Whitten et al., Planet Space Sci. 19:243–250, 1971; Molina-Cuberos et al., J. Geophys. Res. Planets 107:3–1, 2002; Haider et al., J. Geophys. Res. Space Phys., 112(A12), 2007; Haider et al., J. Geophys. Res. Space Phys., 114(A3), (2009); Haider et al., J. Geophys. Res. Space Phys. 121:6870–6888, 2016). Keywords D layer · Modeling · Chemistry

22.1 Mars 4 and 5 Measurements There is no measurement in the daytime lower ionosphere of Mars. The night time ionosphere has been observed from Mars 4 and Mars 5 on 10 and 18 February 1974 respectively (Savich and Samovol 1976; Verigin et al. 1991). These observations were carried out by radio occultation experiment. At the first exit the solar zenith angle at the point of contact of the radio beam with planetary surface was χ ~ 127°, aerographic coordinates of this point were latitude 90° S, longitude 236° W, the local time was about 0330 and the season in this Martian region was autumn. The second exit was in spring at coordinates 38° N, 214° W and local time 0430, with χ ~ 106°. The Fig. 22.1 represents the electron density profiles observed by Mars 4 and Mars 5. It appears that these © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_22

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22 Lower Ionosphere of Mars

Fig. 22.1 Distribution of electron densities in the nighttime ionosphere of Mars. from Mars 4 and Mars 5 on 10 and 18 February 1974 respectively (from Savich and Samovol 1976; Verigin et al. 1991)

measurements contain three peaks at altitude ranges 180–200 km, 110–130 km and 30–35 km, which were produced by different physical processes. The first layer at ~ 180–200 km may exist primarily due to transport of photoelectrons and ions from the dayside ionosphere, which recombined with the night time ionosphere across the terminator (cf. Fox et al. 1993). The second and third layers at altitude ranges 110–130 km and 30–35 km were produced due to precipitation of solar wind electron and GCR respectively (Haider et al. 2007, 2009). The third layer is known as D peak in the lower ionosphere of Mars. D peak occurs due to high efficiency of electron attachment to Ox molecules, which entails that concentration of negative ions is higher than that of electron below ~ 30 km.

22.2 Galactic Cosmic Ray (GCR) Ionization GCR impact ionization is calculated in the lower ionosphere of Mars from energy loss model given by Eq. 10.29 in Chap. 10. Figure 22.2 represents the ion production rates of CO2 + , N2 + , Ar+ , O2 + , CO+ , H2 O+ and H2 + at latitude 67o N and longitude 200o E due to GCR impact at solar zenith 80° (Haider et al. 2008). These calculations

22.3 Chemistry of the Lower Ionosphere

201

Fig. 22.2 Ion production rates of CO2 + , N2 + , Ar+ , O2 + , CO+ , H2 O+ and H2 + at latitude 67° N and longitude 200° E (from Haider et al. 2008)

were made during the summer. In this calculation neutral densities of 12 gases (CO2 , N2 , O2 , O, CO, Ar, O3 , H2 , H2 O, NO, NO2 , and HNO3 ) in the lower atmosphere of Mars were taken from Molina-Cuberos et al. (2002). The production rates of O+ , O3 + , NO+, NO2 + , and HNO3 + are very low, therefore they are not plotted in this figure.

22.3 Chemistry of the Lower Ionosphere Sheel and Haider (2016) have calculated nine positive ions (H3 O+ (H2 O)4 , + + + H3 O+ (H2 O)3 , H3 O+ (H2 O)2 , H3 O+ H2 O, H3 O+ , CO+ 2 , O2 CO2 , NO , and O2 ) and − − − − − − eight negative ions (CO3 (H2 O)2 , CO3 H2 O, CO3 , CO4 , NO2 (H2 O)2 , NO3 H2 O, and NO− 3 (H2 O)2 ) in the lower ionosphere of Mars. In the positive ion chem+ istry 100% ions of O+ 2 and CO2 are produced initially due to impact of GCR. These ions are fully destroyed by H2 O in the formation of H3 O+ . Once H3 O+ has been produced water vapour molecules are attached with it. The ion H3 O+ is then lost by three body reaction in the formation of cluster hydrated ions H3 O+ (H2 O)n for n = 1, 2, 3, and 4.

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22 Lower Ionosphere of Mars

In the negative ion chemistry O− and O− 2 ions are produced initially due to electron − capture process. Later in presence of three body reactions CO− 3 and CO4 were − − − produced due to loss of O and O2 with CO2 , respectively. The ion CO3 is destroyed − − by NO forming NO− 2 . When CO3 and NO2 ions are formed, H2 O is attached with them. The production and loss reactions in the formation of positive and negative cluster ions are given below: H2 O + (H2 O)n + H2 O + M ↔ H3 O + (H2 O)n+1 + M C O3− (H2 O)n + H2 O + M ↔ C O3− (H2 O)n+1 + M N O2− (H2 O)n + H2 O + M ↔ N O2− (H2 O)n+1 + M The chemical reactions are reversible. The hydrated ions are the dominant in lower ionosphere of Mars. These ions are producing the D layer in the lower ionosphere of Mars. The densities of these major ions were reduced by 1–2 orders of magnitude near the surface during the major dust storm as compared to that estimated for absence of dust storm period (Haider et al. 2010, 2015; Sheel and Haider 2012, 2016). In absence of dust storm the chemistry of dust reactions were not considered in the chemical

Fig. 22.3 A schematic diagram illustrating the reaction pathways that are important in the chemistry of D region of Martian ionosphere. The primary source of ionization, galactic cosmic rays that initiate the ion-neutral chemistry in the D region is given at the top. The symbol e− refers to electrons. The neutral species are shown in the circle. The ions taking part in the chemistry are outlined by rectangles (from Haider et al. 2010)

22.4 D Layer Ionosphere of Mars

203

model (Sheel and Haider 2016). The dust reactions are playing very important role in the sink process of ion-dust model (Haider et al. 2010). The Fig. 22.3 represents the schematic diagram of ion-aerosol and ion-neutral reactions in the lower ionosphere of Mars.

22.4 D Layer Ionosphere of Mars GCR has been reported earlier an important source of D layer in the lower ionosphere of Mars at altitude 30 km with electron density ~ 1 × 102 cm−3 (Whitten et al. 1971; Molina-Cuberos et al. 2002; Haider et al. 2009, 2010). Recently, hard X-rays (0.5– 3Å) have also been considered a new source of D region ionosphere of Mars (Shah et al. 2021). The D-peak density produced by hard X-rays is obtained as 8 × 102 cm−3 , which is larger by about an order of magnitude than that estimated by GCR impact ionization. The D layer in presence of X-ray flare is larger by 1–2 orders of magnitude than that produced in absence of flare. Figure 22.4 represents the electron density profiles in presence and absence of X-ray flare. In this figure the electron density profile due to GCR impact ionization is also shown for comparison. The electron density in the D-region is calculated as ne = ∑ni + − ∑ni − , where ∑ni + is the sum of all positive ion densities and ∑ni − is the sum of all negative ion densities. Using charge neutrality and steady state condition from Eqs. 10.2 and 10.7, the electron density ne is obtained by iteration process. The steady state condition 200 SZA~71o

Altitude (km)

150

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Flare (6/4/2001) Flare (17/3/2003) Non-flare (6/4/2001) Non-flare (17/3/2003) GCR

50

0 10 1

10 2

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Fig. 22.4 The flare profiles in these figures are shown by red (star) lines. The estimated non-flare electron density profiles due to impact of X-rays (0.5–90 Å) on 6 April, 2001 and 17 March, 2003 are shown in Fig. 22.4c (see red triangle and blue star). The estimated flare electron density profiles due to impact of X-rays (0.5–90 Å) on 6 April, 2001 and 17 March, 2003 are also shown in this figure (see red and blue lines). The estimated electron density profiles due to GCR impact is shown from dashed green line (from Shah et al. 2021)

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22 Lower Ionosphere of Mars

is valid in the D-region of Mars’ ionosphere because the chemical life time is much lower than the transport time in this region. The chemical reactions used in this chemical model is given in Appendix F and G.

References Fox, J.L., Brannon, J.F., Porter, H.S.: Upper limits to the nightside ionosphere of Mars. Geophys. Res. Lett. 20(13), 1339–1342 (1993) Haider, S.A., Batista, I.S., Abdu, M.A., et al.: Dust storm and electron density in the equatorial D region ionosphere of Mars: Comparison with Earth’s ionosphere from rocket measurements in Brazil. J. Geophys. Res. Space Phys. 120(10), 8968–8977 (2015) Haider, S.A., Batista, I.S., Abdu, M.A., et al.: Flare X-ray photochemistry of the E region ionosphere of Mars. J. Geophys. Res. Space Phys. 121(7), 6870–6888 (2016) Haider, S.A., Singh, V., Choksi, V.R., Maguire, W.C., Verigin, M.I.: Calculated densities of H3 O+ (H2 O)n, NO2 − (H2 O)n, CO3 − (H2 O) n and electron in the nighttime ionosphere of Mars: Impact of solar wind electron and galactic cosmic rays. J. Geophys. Res. Space Phys. 112(A12) (2007) Haider, S.A., Sheel, V., Singh, V., Maguire, W.C., Molina-Cuberos, G.J.: Model calculation of production rates, ion and electron densities in the evening troposphere of Mars at latitudes 67 N and 62 S: seasonal variability. J. Geophys. Res. Space Phys. 113(A8) (2008) Haider, S.A., et al.: D, E, and F layers in the daytime at high-latitude terminator ionosphere of Mars: comparison with Earth’s ionosphere using COSMIC data. J. Geophys. Res. Space Phys. 114(A3) (2009) Haider, S.A., Seth, S.P., Brain, D.A., et al.: Modeling photoelectron transport in the Martian ionosphere at Olympus Mons and Syrtis Major: MGS observations. J. Geophys. Res. Space Phys. 115(A8) (2010) Molina-Cuberos, G.J., Lichtenegger, H., Schwingenschuh, K., et al.: Ion-neutral chemistry model of the lower ionosphere of Mars. J. Geophys. Res. Planets 107(E5), 3–1 (2002) Savich, N.A., Samovol, V.A.: The night-time ionosphere of Mars from Mars 4 and Mars 5 dualfrequency radio occultation measurements. Space Research XVI, 1009–1011 (1976) Shah, S.Y., et al.: A coupled model of the D and E regions of Mars’ ionosphere for flare and non-flare electron density profiles. Icarus 361, 114403 (2021) Sheel, V., Haider, S.A.: Calculated production and loss rates of ions due to impact of galactic cosmic rays in the lower atmosphere of Mars. Planet Space Sci. 63, 94–104 (2012) Sheel, V., Haider, S.A.: Long-term variability of dust optical depths on Mars during MY24–MY32 and their impact on subtropical lower ionosphere: climatology, modeling, and observations. J. Geophys. Res. Space Phys. 121(8), 8038–8054 (2016) Verigin, M.I., Shutte, N.M., Galeev, A.A., et al.: Ions of planetary origin in the Martian magnetosphere (Phobos 2/TAUS experiment). Planet Space Sci. 39(1–2), 131–137 (1991) Whitten, R.C., Poppoff, I.G., Sims, J.S.: The ionosphere of mars below 80 km altitude—I quiescent conditions. Planet Space Sci. 19(2), 243–250 (1971)

Chapter 23

Conductivity

Abstract The ionosphere of Mars has been understood mainly by the observations of ion and electron density profiles. A cavity between lower ionosphere and the surface of Mars provide reflection and attenuation of Electro-magnetic (EM) waves of Extremely low Frequency (ELF) and Very Low Frequency (VLF). These signals produce resonance in the cavity due to conducting effect of ionosphere and surface. The low frequency EM signals depend strongly on ionospheric conductivity. The ionospheric conductivity has been calculated in the lower ionosphere of Mars (Molina-Cuberos et al., Radio Sci. 41:1–8, 2006; Cardnell et al., J. Geophys. Res. 121:2335–2348, 2016; Haider et al., Adv. Space Res. 63:2260–2266, 2019). The conductivity depends on positive and negative ion densities and mobility. Keywords Dusty ionosphere · Dust storm · Aerosol charging

23.1 Ion Conductivity in the Lower Ionosphere The electrical conductivity is not observed in the troposphere of Mars. However, such charging would occur within the clouds observed on Mars. The main source of ionization in the troposphere of Mars is the GCR. The conductivity increases with altitude in the troposphere. Since the ion-neutral collision frequency is very large in the troposphere, the effect of convection will be insignificant and the ions will not be able to move under the influence of the ambient electric field. The conductivity depends on the number density of charged particles according to the relation, σtotal = eμ (

∑

ρ+ +

∑

ρ−)

(23.1)

where σ is the total ion conductivity, μ is the mean ion mobility, e is elementary charge, ρ+ and ρ− are positive and negative ion densities in the lower ionosphere of Mars respectively. Using formula of Michael et al. (2007) the mean ion mobility is calculated to be ~ 0.02 m2 V−1 s−1 . By using this value of the mobility the total ion conductivities has been calculated in the presence and absence of dust storm in the troposphere of Mars. © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_23

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23.2 Effect of Dust Storm on Conductivity Dust cloud particles have been observed in the troposphere of Mars (Montmessin et al. 2006). Haider et al. (2010) have calculated the total ion conductivities in presence and absence of dust aerosols in the troposphere of Mars. They reported that during dust storms the total ion conductivity is reduced by ∼ 2 orders of magnitude near the surface of Mars. Recently, Haider et al. (2019) have calculated the atmospheric conductivities and compared with other model results between 0 and 50 km due to GCR impact ionization in presence and absence of dust storm in the Martian ionosphere which is shown in Fig. 23.1. The blue line with square represent the ion conductivity profile calculated for MY 25 during dust storm at τ = 1.7 and Ls = 210° at low latitude range (25–35° S). The ion conductivity profile given by Molina-Cuberos et al. (2006) in absence of dust storm is represented by pink line with triangle. The red dotted line with circle, blue line with star and brown dotted line represent the ion conductivities estimated by Cardnell et al. (2016) in the presence of high dust, standard dust and low dust at τ = 1.7, 0.5 and 0.1, respectively at Ls = 244°. Haider et al. (2019) reported that the calculated conductivity at peak altitude is lowered by ~ 2 orders of magnitude than that estimated by Molina-Cuberos et al. (2006) in absence of dust. Molina-Cuberos et al. (2006) and Cardnell et al. (2016) did not found clear peak in their modelled conductivities. 50

Altitude (km)

40

30

20

GCR and high dust (present) GCR and No dust (Molina-Cuberos et al., 2006) GCR and High Dust (Cardnell et al., 2016) GCR and Standard Dust (Cardnell et al., 2016) GCR and Low Dust (Cardnell et al., 2016)

10

0 10 -13

10 -12

10 -11

10 -10

10 -9 -1

10 -8

10 -7

-1

Conductivity (Ohm m ) Fig. 23.1 Comparison of ionospheric conductivities between Haider et al. (2019) and other model calculations carried by Molina-Cuberos et al. (2006) and Cardnell et al. (2016) (from Haider et al. 2019)

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23.3 Ion-Dust Model For the first time Haider et al. (2010) have developed an ion-dust aerosol model for the calculation of the conductivity and the densities of positive and negative ions in the Mars’ D region ionosphere. The formation of this model is purely a sequence of algebraic expressions, which yield solutions after sufficient iteration for electron concentration and all the individual ions currently in the model, i.e., positive ions, negative ions, and positively and negatively charged aerosols. Using chemical reactions from Appendixes F and G under photochemical equilibrium conditions, the positive and negative ion concentrations of atmospheric gases and charged aerosol are given below: ∑ dρi+ K i ρk ρi+ − αii ρi+ ρi− − αe ρi+ ρe − [ i ρi+ − α1 Ai− ρi+ − q + = qi+ − A dt k (23.2) ∑ dρi− = qi− − K i ρk ρi− − αii ρi+ ρi− − [ i ρi− − α2 Ai+ ρi− − qA− dt k

(23.3)

d Ai+ = qA+ − βe Ai+ ρe − α Ai+ Ai− − α2 Ai+ ρi− dt

(23.4)

d Ai− + − − + = q− A − α Ai Ai − α1 Ai ρi dt

(23.5)

where ρi+ and ρi− are the positive and negative ion concentrations, respectively, in the presence of aerosols, Ai+ and Ai− are the positive and negative charged aerosol concentrations, respectively, βe and αe are the charged aerosol and ion–electron recombination coefficients, respectively, ρe is the electron concentration, qi+ and qi− are the total ion production rates of positive and negative ions, respectively, q+ A and q− A are the production rates of positive and negative charged aerosols, respectively, α is the charged aerosol-charged aerosol recombination coefficient, α1 is the positive ion and negative-charged aerosol recombination coefficient, α2 is the negative ion and positive charged aerosol recombination coefficient, [ i is the coefficient rate for the loss of ith ion by photon collision, αii is the ion-ion recombination coefficient, Ki is the rate coefficient of the reaction to remove the ion species i by the reaction with neutral concentration ρk of species k. The Eqs. 23.2–23.5 have been used in Eq. 23.1 for the calculation of ion and electron conductivities in presence and absence of dust storm.

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23.4 Chemistry of Dusty Ionosphere The ion dust chemistry consist 106 chemical reactions (see Appendix F and G) which include the ion-neutral reactions, photodissociation of positive and negative ions, electron attachment to neutrals and photodetachment of negative ions, charged aerosol, positive ion-negative charged aerosol, electron—positive charged aerosol, ion-aerosol attachment, and negative ion-positive charged aerosol reactions. The flux of incident GCR is exponentially attenuated in the lower atmosphere and is calculated between values 103 –10−5 particles m−2 s−1 GeV−1 ster−1 at energy range of 1– 1000 GeV (Molina-Cuberos et al. 2002; Haider et al. 2009). The dust reactions are playing very important role in the loss process of hydrated ions (see Eqs. 23.2–23.5).

23.5 Atmospheric Electricity Martian dust is known to impact various photochemical reactions in the Mars’ atmosphere. The ion composition and chemistry of dust present in the Martian troposphere is required to understand the atmospheric electricity (Leblanc et al. 2008a, b; Harrison et al. 2008, 2016; Cardnell et al. 2016; Esposito et al. 2016; Haider et al. 2019). Attachment of ions to dust particles is an important loss process for atmospheric electricity, which may produce lightning in the Martian atmosphere. In presence of dust storms large scale electrostatic fields are generated by charged dust. The positive and negative charged dusts are impacting with the atmospheric gases O2 and O and produce O3 + and O3 − respectively. Later O3 is formed due to dissociative recombination of O3 + with electron and photo dissociative attachment of O3 − respectively. This process is known as tribo-electricity. In presence of triboelectricity lightning can also occur in the Martian atmosphere (Haider et al. 2010). The typical dust abundances can induce 10–50% increase in O3 abundances during the lightning on Mars (Lindner et al. 1988).

23.6 Aerosol Charging The aerosol particles can move in the troposphere due to the influence of gravity field. These aerosol particles can interact with the light ions of atmosphere during their movement. This process produces electric fields several times larger than the electric fields produced by elementary charges during the dust storm period. The production of large electric fields and its associated electric discharges are responsible for the occurrence of lightning during the dust storm period on Mars (Tolendo-Redondo et al. 2017). The lightning discharge may be produced within dust storms. It is known that dust devils or storms occur on Mars and it is expected that lightning may be present within the dust devils on Mars. The heavy loss of ions occurred in presence of high

References

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dust. This is the evident from Fig. 23.1 where the surface conductivity is reduced by ~ 2 orders of magnitude with increasing dust opacity. The conductivity in the presence of high dust is calculated to be ~ 10–13 Ω−1 m−1 at the surface of Mars.

References Cardnell, S., Witasse, O., Molina-Cuberos, G.J., et al.: A photochemical model of the dust loaded ionosphere of Mars. J. Geophys. Res. 121, 2335–2348 (2016) Esposito, F., Molinaro, R., Popa, C.I., et al.: The role of the atmospheric electric field in the dustlifting process. Geophys. Res. Lett. 43(10), 5501–5508 (2016) Haider, S.A., Pabari, J.P., Masoom, J., Shah, S.Y.: Schumann resonance frequency and conductivity in the nighttime ionosphere of Mars: a source for lightning. Adv. Space Res. 63(7), 2260–2266 (2019) Haider, S.A., et al.: D, E, and F layers in the daytime at high-latitude terminator ionosphere of Mars: comparison with Earth’s ionosphere using COSMIC data. J. Geophys. Res. Space Phys. 114(A3) (2009) Haider, S.A., Seth, S.P., Brain, D.A., et al.: Modeling photoelectron transport in the Martian ionosphere at Olympus Mons and Syrtis Major: MGS observations. J. Geophys. Res. Space Phys. 115(A8) (2010) Harrison, R.G., Barth, E., Esposito, F., et al.: Applications of electrified dust and dust devil electrodynamics to Martian atmospheric electricity. Space Sci. Rev. 203(1), 299–345 (2016) Harrison, R.G., Aplin, K.L., Leblanc, F., Yair, Y.: Planetary atmospheric electricity. Space Sci. Rev. 137(1), 5–10 (2008) Leblanc, F., Aplin, K., Yair, Y., et al., Planetary atmospheric electricity, vol. 30. Springer Science and Business Media (2008a) Leblanc, F., Witasse, O., Lilensten, J., et al.: Observations of aurorae by SPICAM ultraviolet spectrograph on board Mars Express: simultaneous ASPERA-3 and MARSIS measurements. J. Geophys. Res. Space Phys. 113(A8) (2008b) Lindner, B.L.: Ozone on Mars: the effects of clouds and airborne dust. Planet Space Sci. 36(2), 125–144 (1988) Michael, M., Barani, M., Tripathi, S.N.: Numerical predictions of aerosol charging and electrical conductivity of the lower atmosphere of Mars. Geophys. Res. Lett. 34(1–5), L04201 (2007) Molina-Cuberos, G.J., Lichtenegger, H., Schwingenschuh, K., et al.: Ion-neutral chemistry model of the lower ionosphere of Mars. J. Geophys. Res. Planets 107(E5), 3–1 (2002) Molina-Cuberos, G.J., Morente, J.A., Besser, B.P., et al.: Schumann resonances as a tool to study the lower ionospheric structure of Mars. Radio Sci. 41, 1–8, RS1003 (2006) Montmessin, F., Bertaux, J.L., Quémerais, E., Korablev, O., Rannou, P., Forget, F., Dimarellis, E.: Subvisible CO2 ice clouds detected in the mesosphere of Mars. Icarus 183(2), 403–410 (2016) Toledo-Redondo, S., Salinas, A., Portí, J., et al.: Schumann resonances at Mars: effects of the daynight asymmetry and the dust-loaded ionosphere. Geophys. Res. Lett. 44(2), 648–656 (2017)

Chapter 24

Dust Storms in the Lower Atmosphere of Mars

Abstract Dust storms have been frequently observed on Mars. While others dust storms are global some are regional and local. Many observations have been carried out between MY10 and MY34 (Martin, Icarus 66:2–21, 1986, 1995; Heavens et al., J. Geophys. Res. Planets 119:1748–1774, 2014; Montmessin et al., Icarus 297:195– 216, 2017; Willame et al., Planet Space Sci. 142:9–25, 2017; Smith, J. Geophys. Res. Planets 124:2929–2944, 2019; Guzewich et al., Geophys. Res. Lett. 46:71–79, 2019; Guzewich et al., J Geophys. Res. Planets 126:e2021JE006825, 2021). TES onboard MGS have observed a GDS in MY 25 at latitude 25° S, where the IR optical depths were increased up to ~ 1.7 at Ls = 210° (Smith, Icarus 167:148–165, 2004; Sheel and Haider, J. Geophys. Res. Space Phys. 121:8038–8054, 2016) (see Fig. 26.1). Later THEMIS onboard Mars Odyssey observed two major dust storms in MY28 and MY 34 at latitude 25°S, where the IR optical depths were increased up to ~ 1.2 and 1.3 at Ls =280° and Ls = 195° respectively (Smith, Icarus 202:444–452, 2009; Smith, J. Geophys. Res. Planets 124:2929–2944, 2019; Montabone et al., Icarus 251:65–95, 2015; Sheel and Haider, J. Geophys. Res. Space Phys. 121:8038–8054, 2016; Smith, J. Geophys. Res. Planets 124:2929–2944, 2019) (see Fig. 26.2a, b). Keywords Dust layers · Optical depth · Dust density

24.1 Infrared Dust Optical Depth: MGS Measurements The earlier measurements of Martian atmosphere by TES on board MGS have provided an insight about the seasonal and annual variability of dust cycle on Mars. The Fig. 24.1 shows the seasonal and latitudinal dependence of dust optical depth as observed by TES between MY 24 and MY 26. During MY 24 a regional dust storm occurred at Ls ~ 225°–245°, where the optical depth was measured as 0.5. During MY 25 a GDS was observed at Ls = 185°–210°, where the dust optical depth increased up to 1.2 (Smith et al. 2002; Sheel and Haider 2016).

© Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_24

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Fig. 24.1 Zonal averaged dust optical depth as a function of Ls and latitude, retrieved from TES for MY 24 to MY26 at 1075 cm−1 (from Smith 2004)

24.2 Infrared Dust Optical Depth: Mars Odyssey Observations THEMIS onboard Mars Odyssey carried out a long-term monitoring of the Mars atmosphere from MY 26 to MY 34 (Smith 2009, 2019). The Figs. 24.2a and 24.2b represent the seasonal variability of zonally averaged dust optical depths from MY 26 to MY 29 and MY 33 to MY 34 respectively (Smith 2009, 2019). The global dust storms on Mars took about 3–4 months for the dust to settle down to the background dust loading. Many of these major dust storms begin in subtropical region in the southern summer season after perihelion, when Mars receives a higher amount of radiative energy from the sun compared to other seasons. The increased solar insolation increases the temperature and wind speed that lift dust from the surface into the atmosphere of Mars (Lemmon et al. 2015; Sheel and Haider 2016). The dust opacity exceeded from 1.0 over most of the planet for several weeks during global dust storms (Sheel and Haider 2016). The zonal- mean UV dust opacity was also observed in the southern hemisphere by SPICAM onboard MEX between MY27 and MY30 (Haider et al. 2019 and Montmessin et al. 2017). It was found the dust optical depth ~ 3 on the surface of Mars in MY 28. Recently, the MSL Curiosity rover has also observed a GDS from Mastcam instrument in MY34 (Guzewich et al. 2019, 2021). It measured the dust optical depth ~ 8.5 at a wavelength of 880 nm. This is a large optical depth because it was measured at a visible wavelength. The GDS MY34 started in mid-May 2018 at Ls ~ 181° in the northern hemisphere and expanded out to the southern hemisphere during June 2018 at Ls ~ 190°. This dust storm was also observed by MCS onboard MRO (Kass et al. 2020). The effect of this dust storm is also observed on water and ozone by NOMAD instrument onboard ExoMars TGO (Stcherbinine et al. 2020; Vandaele et al. 2019).

24.2 Infrared Dust Optical Depth: Mars Odyssey Observations

213

Fig. 24.2 a Zonally averaged dust optical depth at 9.3 μm band as a function of Ls and latitude, retrieved from THEMIS for MY 26 to MY 29 (from Smith 2009), b shows Zonally averaged dust optical depth as a function of Ls and latitude, retrieved from THEMIS for MY 33 and MY 34 (from Smith 2019)

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24.3 Dust Layers in the Martian Atmosphere The high-altitude layers of aerosols in Mars’ atmosphere have been imaged since 1978 Mariner 9 (Anderson and Leovy 1978). Similar observations were made from Viking orbiters (Jaquin et al. 1986), which suggested that the dust can reach up to ~ 50 km in the Martian atmosphere. The TES and MCS instruments have also observed a layer/peak of dust mixing ratio at altitude ~ 60 km (Clancy et al. 2010). These layers are also known as “detached dust layers”, which were observed throughout the year at tropical and sub-tropical latitudes (Heavens et al. 2014; Guzewich et al. 2014). The dust layers have different sizes (Haider et al. 2015) and thus plays a key role in the climate modeling of Mars.

24.4 Characteristics of Dust and Its Size Distributions Since several decades, dust has been considered an important parameter by controlling the climate of Mars (Gierasch and Goody 1972). It can absorb and scatter solar radiation, and its subsequent heating effect on the atmosphere is significant (Pollack et al. 1990).The size of dust particles determines radiative effect and time of sedimentation due to gravity and their drag (Kahre et al. 2017). The smaller dust particles may stay longer in the atmosphere compared with the larger particles and can be carried farther away by the atmospheric circulation. As a result, the vertical distribution of dust mixing ratio will be different for different sizes of dust particles, and so leading to different heating effect on the atmosphere. The particle sizes of dust, their distribution and the effective radius in the Martian atmosphere are uncertain. Smith et al. (2013) used fixed effective radius 1.5 μm in the Martian atmosphere. Guzewich et al. (2014) reported that effective radius is always 1.0 μm throughout the Martian year below 40 km. Recently, dust aerosol size has been determined using direct Sun imaging, passive sky spectroscopy and ultraviolet sensor systematic measurements from MSL (Vicente-Retortillo et al. 2020; Chen-Chen et al. 2019). To study the triboelectrification of the Mars’ dust caused by erosion in a storm, the experiments should be developed to characterize the particle size and electrostatic charge distributions.

24.5 Effects of Dust in the D Region Ionosphere There is no observation for the D region ionosphere of Mars. Several theoretical models have confirmed that Mars has a permanent D region in the lower ionosphere (Whitten et al. 1971; Molina-Cuberos et al. 2001, 2002; Haider et al. 2007, 2008, 2009; Shah et al. 2021). Dust cloud particles have been observed in the troposphere of Mars (Montmessin et al. 2006). The electrical conductivity is not observed in

24.6 Density Distribution Model of Aerosol Particles

215

the Martian troposphere. However, such charging would occur within the clouds observed on Mars. Haider et al. (2010) have developed a photochemical model for the calculation of ion concentration and conductivity in the D region of the Martian ionosphere in presence and absence of a dust storm. Knowledge of the ion composition and chemistry in presence of dust are required to fully understand the atmospheric electricity and wave propagations in the Martian atmosphere. In the dust storm, the total ion conductivity in the troposphere is reduced by an order of magnitude (Haider et al. (2010) (see Fig. 26.3). The concentrations of water cluster ions H+ (H2 O)n, NO2 − (H2 O)n, and CO3 − (H2 O)n is also decreased by about 2 orders of magnitude near the surface of Mars if dust aerosols are present. These results indicate that during massive dust storms the lower ionosphere of Mars is significantly perturbed and a hole in the ion concentrations may appear at the bottom of the D layer until this anomalous situation returns to the normal condition after the storms.

24.6 Density Distribution Model of Aerosol Particles Haider et al. (2015) have developed the dust aerosol model between altitudes 0 km and 80 km for different size of dust aerosol particles. This model calculates ion densities at different latitudes, altitudes, and solar longitudes. The model includes chemical processes like ion-neutral collisions, electron-neutral collisions, dissociation of positive and negative ions, electron detachment of anions, ion-dust attachment, and ion-ion and ion–electron recombination. This model is used to estimate the densities of positive and negative ions due to precipitation of GCR in presence and absence of dust storms at solar zenith angle 60°. Initially, GCR ionizes CO2 , N2 , Ar, O2 , CO, H2 , H2 O, O, O3 , NO, NO2 , and HNO3 , and production rates of 12 ions were obtained in the lower ionosphere. Later these production rates were used in the calculations of positive and negative ion densities. This figure represents the concentration profiles of dust aerosol of size 0.2, 0.6, 1.0, 1.4, 2.0, and 3.0 μm in spring season of MY25 at latitude 10° S and Ls = 200°, 220°, 250°, and 280° for high dust storm (τ = 1.9), medium dust storm (τ = 1.0), low dust storm (τ = 0.5), and the absence of dust storm (τ = 0.1), respectively. The concentrations of dust aerosols of size > 3 μm are nearly zero in the lower atmosphere of Mars. The background optical depth is equal to 0.1 when the atmosphere is clean from the dust loading. Thus, aerosol particles of τ = 0.1 are present in the clean atmosphere during the absence of dust storm. It is found that the dust density is high for smaller sized particles. Since the gravitation settling velocity is directly proportional to the size of dust particles, the larger particles settle down quickly and their concentrations are lower than the small particles. Based on MGS measurements in MY25, the low dust storm at Ls = 250° for τ = 0.5 is confined to altitude region

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Fig. 24.3 Altitude profiles of dust densities at 10°S for different size of aerosol particles in the spring season of MY25 at a Ls = 200°, b Ls = 220°, c Ls = 250°, and d Ls = 280° (from Haider et al. 2015)

References

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~ 25–30 km, while medium and high dust storms at L = 220° and Ls = 200° for τ = 1.0 and 1.9, respectively, have risen to altitude ~ 50 km (see Fig. 24.3a–c). The dust storm is transported from the surface to about 80 km through dynamical coupling processes providing expansion of the lower atmosphere, which increased the densities and temperature during this event. Guzewich et al. (2014) and Heavens et al. (2014) have studied vertical distribution of dust aerosol for the same dust storm conditions in the dayside tropical atmosphere of Mars. They have found two distinct layers in the dust profiles, one at altitude ~ 20–30 km and other at altitude ~ 45–65 km from the observations by CRISM and MCS experiments, respectively.

References Anderson, E., Leovy, C.: Mariner 9 television limb observations of dust and ice hazes on Mars. J. Atmos. Sci. 35(4), 723–734 (1978) Chen-Chen, H., Pérez-Hoyos, S., Sánchez-Lavega, A.: Dust particle size and optical depth on Mars retrieved by the MSL navigation cameras. Icarus 319, 43–57 (2019) Clancy, R.T., Wolff, M.J., Whitney, B.A., et al.: Extension of atmospheric dust loading to high altitudes during the 2001 Mars dust storm: MGS TES limb observations. Icarus 207(1), 98–109 (2010) Gierasch, P.J., Goody, R.M.: The effect of dust on the temperature of the Martian atmosphere. J. Atmos. Sci. 29(2), 400–402 (1972) Guzewich, S.D., Smit, M.D., Wolff, M.J.: The vertical distribution of Martian aerosol particle size. J. Geophys. Res. Planets 119(12), 2694–2708 (2014) Guzewich, S.D., Lemmon, M., Smith, C.L., et al.: Mars Science Laboratory observations of the 2018/Mars year 34 global dust storm. Geophys. Res. Lett. 46(1), 71–79 (2019) Guzewich, S.D., Way, M.J., Aleinov, I., et al.: 3D Simulations of the early martian hydrological cycle mediated by a H2 –CO2 Greenhouse. J. Geophys. Res. Planets. 126(7), e2021JE006825 (2021) Haider, S.A., Batista, I.S., Abdu, M.A., et al.: Dust storm and electron density in the equatorial D region ionosphere of Mars: Comparison with Earth’s ionosphere from rocket measurements in Brazil. J. Geophys. Res. Space Phys. 120(10), 8968–8977 (2015) Haider, S.A., Pabari, J.P., Masoom, J., Shah, S.Y.: Schumann resonance frequency and conductivity in the nighttime ionosphere of Mars: a source for lightning. Adv. Space Res. 63(7), 2260–2266 (2019) Haider, S.A., Singh, V., Choksi, V.R., Maguire, W.C., Verigin, M.I.: Calculated densities of H3 O+ (H2 O)n, NO2 − (H2 O)n, CO3 − (H2 O) n and electron in the nighttime ionosphere of Mars: impact of solar wind electron and galactic cosmic rays. J. Geophys. Res. Space Phys. 112(A12) (2007) Haider, S.A., Sheel, V., Singh, V, Maguire, W.C., Molina-Cuberos, G.J.: Model calculation of production rates, ion and electron densities in the evening troposphere of Mars at latitudes 67 N and 62 S: Seasonal variability. J. Geophys. Res. Space Phys. 113(A8) (2008) Haider, S.A., Abdu, M.A., Batista, I.S., et al.: D, E, and F layers in the daytime at high-latitude terminator ionosphere of Mars: comparison with Earth’s ionosphere using COSMIC data. J. Geophys. Res. Space Phys. 114(A3) (2009) Haider, S.A., Seth, S.P., Brain, D.A., et al.: Modeling photoelectron transport in the Martian ionosphere at Olympus Mons and Syrtis Major: MGS observations. J. Geophys. Res. Space Phys. 115(A8) (2010) Heavens, N.G., Johnson, M.S., Abdou, W.A., et al.: Seasonal and diurnal variability of detached dust layers in the tropical Martian atmosphere. J. Geophys. Res. Planets 119(8), 1748–1774 (2014)

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Chapter 25

Lightning on Mars

Abstract It is well known that lightning occurs frequently within water clouds on Earth (Berger and Vogelsanger in Planet Electrodyn 1:489–510, 1969, Farrell et al. in J Geophys Res Planet 109(E3), 2004; Srinivasan and Gu in 2006 Canadian conference on electrical and computer engineering 2006 May 7, pp 2258–2261. IEEE, 2006). The lightning may also occur within the dust storm or dust devils due to charge transfer by triboelectricity among the dust particles (Farrell et al. in J Geophys Res Planet 109(E3), 2004). The lightning discharge emits low frequency electromagnetic waves along with visible light. The low frequency waves are reflected from the Martian ionosphere and can give rise to Schumann Resonance (SR) in the cavity formed by the Martian surface and the ionosphere. The SR frequency strongly depends on atmospheric conductivity, lightning, the ionospheric turbulence and the surface activity of Mars (Barr et al. in J Atmospher Solar-Terr Phys 62:1689–1718, 2000; Molina-Cuberos et al. J Geophys Res Planets 107:3–1, 2002). There are a few approaches like finite differences (Yang and Pasko in Geophys Res Lett 32(3), 2005), two dimensional telegraph equations (Pechony and Price in Radio Sci 39:1, 2004), full wave equation (Galuk et al. in Telecommun Radio Engin 74(15), 2015) and three dimension Transmission Line Matrix (TLM) method (Toledo-Redondo et al. in Geophys Res Lett 44:648–656, 2017) to study the SR frequency in the planetary atmospheres. The atmosphere of Mars is a non-homogeneous medium, where the low frequency waves can oscillate (Morente et al. in J Geophys Res Space Phys 109(A5), 2004) and the densities are changing with height. In homogeneous atmosphere the gaseous composition is constant with altitude. Keywords Lightning · Dust discharge · Schumann Resonance

25.1 SR Model for Lightning on Mars Haider et al. (2019) solved Maxwellian equations of electromagnetic waves in the non-homogeneous medium to derive the SR model within the Martian atmosphere cavity formed in the lower ionosphere between 0 and 70 km. Using the Maxwellian formula of electric and magnetic fields, SR frequencies are calculated in presence of © Springer Nature Singapore Pte Ltd. 2023 S. A. Haider, Aeronomy of Mars, Astrophysics and Space Science Library 469, https://doi.org/10.1007/978-981-99-3138-5_25

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25 Lightning on Mars

dust storm. They considered a wave of order e−iωt in spherical coordinates, r, θ, and φ. The magnetic field wave is defined as B = B0 e−iωt

(25.1)

∂B = −i ω B ∂t

(25.2)

and its time derivative is

The Maxwell equations become now ∇ ×E=−

∂B = i ωB ∂t

(25.3)

∂E ∂t

(25.4)

and ∇ × B = μ0 J + μ0 ε0

where B is the magnetic field of amplitude B0 , E is the electric field, t is the time, ω is the resonance frequency, J is the current density, μ0 and ε0 are permeability and permittivity in space respectively. Transverse Magnetic (TM) wave propagation is considered in this model and it is assumed that the fields are independent of the azimuthal angle ϕ. The Maxwell equations specify that TM modes with no ϕ dependence involve only E r , E θ and Bϕ components. The two curl equations can be combined after assuming a time dependence of e−iwt , into w2 B − (∇ × ∇ × B) = 0 c2

(25.5)

] w2 B [ − ∇(∇.B) − ∇ 2 B 2 c

(25.6)

which is simplified as

Here, ∇.B = 0 because magnetic field lines are in closed loop and its divergence does not exist. The ϕ component of Eq. (25.6) is written as [ ] ) ) 1 ∂ ( ω2 1 ∂ ∂2 ( sin θ r Bϕ = 0 r Bϕ + 2 (r Bϕ ) + c2 ∂ r2 r ∂θ sin θ ∂θ

(25.7)

The angular part of Eq. (25.7) can be transformed into ( )) ( [ ] ) ∂ r Bφ 1 ∂ ( ∂ 1 ∂ r Bφ sin θ r Bϕ = sin θ − ∂θ sin θ ∂θ sin θ ∂θ ∂θ sin2 θ

(25.8)

25.1 SR Model for Lightning on Mars

223

Substituting Eq. (25.8) into Eq. (25.7), we get [ ] ( ) ) ) 1 ∂ ∂ ( ω2 1 r Bφ ∂2 ( =0 sin θ r Bφ − r Bϕ + 2 (r Bϕ ) + c2 ∂ r2 r sin θ ∂θ ∂θ sin2 θ

(25.9)

We have solved above equations using Legendre Polynomials Plm (cos θ ) with m = ± 1. Now, using Bφ (r, θ ) =

u l (r ) m Pl (cos θ ) r

(25.10)

into Eq. (25.9), we get differential equation for ul (r) as given below. ω2 u l (r ) m u l (r ) ∂ m ∂ 2 u l (r ) m P + cot θ 2 θ + P θ p (cos θ ) (cos ) (cos ) l l c2 ∂ r2 r ∂θ l u l (r ) ∂ 2 u l (r ) m P (cos θ ) = 0 + 2 2 Plm (cos θ ) − 2 r ∂ θ r sin θ l

(25.11)

where l = 1, 2, 3 … are modes of SR. In TM modes, m does not vary and it is not a relevant quantity (Jackson 1999). Therefore, associated Legendre Polynomial Plm (cos θ ) can be written as P11 (cos θ ) = −sin θ;P21 (cos θ ) = −3 cos θ sin θ and P31 (cos θ ) = −1.5 (5 cos2 θ − 1) sin θ. The maximum wave amplitude θ = 90° is taken for the transverse wave. To solve this equation lower height of cavity as r = Rm while the upper height of cavity as h