By Brita K. Irving
R E C O M M E N D E D :
Dean, College o f Natural Science and M athem atics Chair, Departm ent o f A tm ospheric Sciences
A P P R O V E D :
A
THESIS
Presented to the Faculty
o f the U niversity o f A laska Fairbanks
in Partial Fulfillm ent o f the Requirem ents
for the D egree o f
M A STE R OF SCIENCE
By
B rita K. Irving, B.S.
Fairbanks, A laska
A ugust 2012
Abstract
R ayleigh lidar observations at Poker F lat Research Range, Chatanika, A laska (65°N, 213°E), have yielded density and tem perature m easurem ents from 4 0 -8 0 km. These m easurem ents have been m ade under clear nighttim e skies since N ovem ber 1997. This thesis presents a study o f M esospheric Inversion L ayers (M ILs) and lidar perform ance at Chatanika. M ILs are identified and characterized in the 4 0 -7 0 km altitude region on 55 o f the 149 w intertim e observations over tw o periods, N ovem ber 1997 - April 2005 and N ovem ber 2007 - M arch 2009, using a new detection algorithm. Investigation o f the M ILs com pared w ith planetary w ave activity as observed by satellite finds a strong correlation betw een the presence of M ILs and the structure o f the planetary waves. These tw o periods are m arked by strong planetary w ave activity and sudden stratospheric w arm ing events. M ILs are found to occur m ore frequently than previously reported at A rctic sites, but less frequently than at low er latitudes. In spring 2012 the existing lidar system w as extended by incorporating a larger aperture telescope and higher pow er laser and field trials w ere conducted. The results from these field trails are presented and the ability o f the new lidar system to extend the scope o f future studies at C hatanika is assessed.
Signature p ag e... i Title P a g e ... ii A bstract ... iii Table o f C o n ten ts... iv L ist o f F ig u re s... vi L ist o f T a b le s ... x L ist o f A p p en d ices...xii
A cknow ledgem ents... xiii
D ed icatio n... xv
Chapter 1. Introduction...1
1.1. The A rctic m iddle atm osphere... 1
1.2. The w ave-driven circulation o f the m iddle atm o sp h e re ... 4
1.3. R ayleigh lid a r...12
1.4. Scope o f this stu d y ... 14
Chapter 2. Principles and Techniques of Rayleigh L idar...15
2.1. N IC T R ayleigh li d a r ...15
2.2. The lidar eq u a tio n ... 19
2.3. L id ar density and tem perature retrievals...27
2.4. Expectations o f the upgraded lidar system ... 32
2.5. S um m ary... 35
Chapter 3. Performance of the Extended Rayleigh Lidar System ... 36
3.1. Introduction...36 Table of Contents
3.2. Extending the R ayleigh lidar s y ste m ...37
3.3. Pulse P ileu p ...41
3.4. Telescope T ria ls...47
3.4.1. 18-19 February 2 0 1 2 ...47
3.4.2. 22 -23 M arch 2012 ... 51
3.4.3. 2 3 -2 4 April and 2 4-2 5 April 2 0 1 2 ...55
3.5. L aser Trials ...61
3.5.1. 2 8 -2 9 M arch 2 0 1 2 ... 61
3.5.2. 3 -4 April 2012 ... 64
3.6. S um m ary... 69
Chapter 4. Mesospheric Inversion Layers at Chatanika and their Relationship to Planetary Wave Structure... 73
4.1. Introduction...73
4.2. M ethods...77
4.2.1. N IC T R ayleigh lidar M IL detection m e tric ...77
4.2.2. TIM ED -SA B ER instrum ent... 83
4.3. M IL s at C hatanika... 84
4.4. M ILs and their relationship to planetary w av es... 92
4.4.1. R eview o f Salby et al. [2 0 0 2 ]...94
4.4.2. SABER observations at 6 5 ° N ...96
4.5. D iscu ssio n ... 104
4.6. S um m ary...107
Chapter 5. Conclusions and Recommendations for Future W ork...109
Figure 1.1. Tem perature o f the E arth ’s atm osphere as a function o f height from the M SISE-90 m odel for 21 June and 22 D ecem ber 2011 at (65°N, 2 1 3 ° E )... 1 Figure 1.2. Schem atic illustration o f January-February stratospheric circulation and eddy heat fluxes (adapted from N ew m an et al. [2 0 0 1 ])... 6 Figure 1.3. Schem atic illustration o f the solstice season pole-to-pole w ave-driven transport circulation in the m esosphere (adapted from H olton and A lexander [2 0 0 0 ])... 7 Figure 1.4. The zonal m ean (a) tem peratures at 70°N w ith dashed line indicating start o f a SSW, (b) wind, (c) EP flux divergence, and (d) gravity w ave forcing averaged for 5 5 - 70°N for (left) model year 1996/1997 and (right) model year 1973/1974 plotted as a function o f altitude and d a y ... 9 Figure 2.1. Schem atic diagram o f the N IC T R ayleigh lidar system at PFRR, Chatanika, A laska (65°N, 2 1 3 ° E )...17 Figure 2.2. Photon count profile plotted as a function o f altitude m easured by the N ICT R ayleigh lidar at P FR R on the night o f 2 -3 January 201 2... 20 Figure 2.3. Signal levels per-laser-pulse as a function o f set num ber on 2 -3 January 2012 (2127-0913 L S T )... 22 Figure 2.4. R elative photon counting error (%) plotted as a function o f altitude from 2 -3 January 20 12... 23 Figure 2.5. D ensity profile plotted as a function o f altitude retrieved from the total photon count profile m easured on 2 -3 January 2012 by the N IC T R ayleigh lidar and norm alized to one at 40 k m ... 28 Figure 2.6. The tem perature profile plotted as a function o f height, m easured by the N ICT R ayleigh lidar on 2 -3 January 2012 (2127-0913 L S T )...30 Figure 3.1. Schem atic diagram o f the Extended R ayleigh lidar system at PFRR, Chatanika, A laska (65°N, 213°E )... 40
List of Figures
Figure 3.2. Schem atic illustration o f pulse pileup behavior in paralyzable and nonparalyzable PM Ts (adapted from Evans [1 9 5 5 ])... 44 Figure 3.3. V ariation o f observed count rate CobS w ith true count rate CTrue for a detector w ith a one second dead tim e... 46 Figure 3.4. Statistical characteristics o f pulse pileup as a function o f the product o f true (or observed) count rate and detector dead time, x = CTRUETd... 46 Figure 3.5. (a) Total lidar signal and (b) photon counting error (%) profiles plotted as a function o f altitude on 18-19 February 2012 (2245-0654 L S T )... 48 Figure 3.6. Signal levels per-laser-pulse plotted as a function o f set num ber on 18-19 February 2012 (2245-0654 L S T )... 49 Figure 3.7. Ratio o f lidar signal (41-inch/24-inch) on 18-19 February 2012 plotted as a function o f altitude (gray line)... 49 Figure 3.8. (a) Total lidar signal and (b) photon counting error (%) profiles plotted as a function o f altitude on 22 -23 M arch 2012 L S T ... 53 Figure 3.9. Signal levels per-laser-pulse plotted as a function o f set num ber on 22-23 M arch 2012 (2039-0529 L S T )... 54 Figure 3.10. R atio o f lidar signal (41-inch/41-inch-ra) on 2 2-23 M arch 2012 plotted as a function o f altitude (gray line)... 54 Figure 3.11. Signal levels per-laser-pulse as a function o f set num ber for April 2008 and 2012...59 Figure 3.12. (a) Photon count profiles and (b) photon counting error (% ) plotted as a function o f altitude for April 2008 and 20 12... 60 Figure 3.13. (a) R atio o f pow er com pensated lidar signals for 24 April 2012 (dark green line) and 25 April 2012 (gray line), and (b) the ratio o f pow er com pensated April 2012 lidar sig n als...60 Figure 3.14. (a) Total lidar signal and (b) photon counting error (%) profiles plotted as a function o f altitude on 2 8 -2 9 M arch 2012 L S T ... 63 Figure 3.15. Signal levels per-laser-pulse plotted as a function o f set num ber on 2 8 -2 9 M arch 2012 (2215-0442 L S T )... 64
Figure 3.16. R atio o f lidar signal (PL9030 41-inch/PL8020 24-inch) on 2 8 -2 9 M arch 2012 plotted as a function o f altitude (gray lin e)... 64 Figure 3.17. Schem atic illustration o f real tim e data set acquisition and processing m ethod on the night o f 4 April 20 12... 65 Figure 3.18. (a) Total lidar signal and (b) photon counting error (%) as a function o f altitude on 4 April 201 2... 66 Figure 3.19. (a) Signal levels and (b) ratio o f lidar signal (PL9030/PL8020) on 4 April
2012 66
Figure 3.20. Tem peratures m easured by the PL8020 and PL9030 on 4 April 2 0 1 2 ...68 Figure 4.1. M onthly distribution o f N IC T Rayleigh lidar observations at C h atan ik a 78 Figure 4.2. (a) Total tem perature profiles as a function o f altitude m easured by the N ICT R ayleigh lidar on (a) 2 5 -2 6 January 2003 LST (b) 2 7 -2 7 January 2005, and (c) 3 February 2008 (thin solid lin e)... 81 Figure 4.3. Sequential 2 h tem perature profiles on (a) 2 5 -2 6 January 2003, (b) 27-2 8 January 2005, and (c) 3 February 200 8... 87 Figure 4.4. Characteristics o f M ILs identified in the average and 2 h tem perature profiles m easured by the N IC T R ayleigh lidar at C h atan ik a... 88 Figure 4.5. M IL phase progression calculated for M ILs identified as significant in tw o or m ore 2 h tem perature profiles... 90 Figure 4.6. (a) Tem perature (K) and (b) w ave geopotential (m2 s-2) at 44°N, as a cross section o f longitude and height, observed by U A RS on 25 D ecem ber 1991... 95 Figure 4.7. (Left) A verage tem perature (K) and (right) geopotential height perturbation (km) at 65°N as a cross section o f longitude and altitude observed by SABER on 26 January 2003 (top), 28 January 2005 (m iddle) and 4 February 2008 U T (bottom ) 98 Figure 4.8. SABER planetary w ave-one (left) and w ave-tw o (right) geopotential am plitudes at 65°N ... 101 Figure 4.9. SABER EP flux divergence at 6 5 ° N ... 102 Figure B.1. Geom etry o f the SAB ER limb approach w ith a tangent height H 0 (adapted from Russell et al. [1999])... 140
Figure B.2. (a) Limb view ing w eighting function for an ideal instrum ent w ith a spectral band o f 585-705 cm -1 [Liou, 1980] and (b) cross sectional view o f the SABER instrum ent [Smith, 2 010]... 142 Figure B.3. V ibrational levels o f triatom ic m olecules such as CO2 (adapted from Petty, [2006])... 144 Figure B.4. M ain CO2 vibrational state energies [G arcia-Com as et al., 20 08]... 144 Figure B.5. Fractional contribution o f CO2 vibrational bands to SABER channel-1 sim ulated radiance for a typical m idlatitude profile (adapted from G arcia-C om as et al. [2008])... 145 Figure B .6 . Sim plified T(p) retrieval (adapted from R em sberg et al. [2008])... 147 Figure B.7. Sim plified Tk retrieval (adapted from G arcia-C om as et al. [2008])... 150 Figure B .8 . SABER v1.07 kinetic tem peratures for equinox (18 M arch 2004) and solstice (15 July 2004) [G arcia-Com as et al., 20 08]... 151 Figure B.9. SABER vibrational tem peratures o f the main CO2 u 2 vibrational states contributing to the 15-p.m channel typical o f (a) m idlatitudes, (b) polar summer, and (c) polar w inter [G arcia-Com as et al., 2008]... 152 Figure B.10. (a) SABER V 1.07 T(z) profile (red curve) com pared w ith a Rayleigh lidar sounding (blue curve) at Table M ountain, California, for 8 June 2002 (adapted from R em sberg et al. [2008])... 155 Figure B.11. Profile o f the average tem perature differences, SABER minus R ayleigh lidar [Rem sberg et al., 2008]...156 Figure B.12. C om parison o f SABER tem perature profile for July w ith a falling sphere clim atology for 15 July at (a) 69°N and (b) o f 70°N [Rem sberg et al., 20 08]... 158
Table 2.1. N IC T R ayleigh lidar system specifications...16
Table 2.2. Lidar signal statistics for 2 -3 January 2012 (2127-0913 L S T )... 24
Table 2.3. Tem perature and density statistics for 2 -3 January 2012 (2127-0913 LST).31 Table 2.4. Tem perature and density statistics for 2 -3 January 2012 (2353-0202 LST). 32 Table 2.5. Lidar signal statistics for 2 -3 January 2012 (2353-0202 L S T )... 33
Table 2.6. Expected lidar signal statistics for 2 h m easurem ents w ith individual system u p g ra d e s...34
Table 2.7. Expected lidar signal statistics for 2 h m easurem ent2 w ith all upgraded system param eters...35
Table 3.1. Extended Rayleigh lidar system specifications...39
Table 3.2. R ayleigh lidar acquisition param eters for 18-19 February 2 0 1 2 ...48
Table 3.3. R ayleigh lidar acquisition param eters for 22 -2 3 M arch 2 0 1 2 ... 53
Table 3.4. R ayleigh lidar acquisition param eters in late April 2008 and 2 0 1 2 ... 57
Table 3.5. Com parison o f pow er com pensated lidar signals in April 2008 and 2 0 1 2 ...58
Table 3.6. R ayleigh lidar acquisition param eters for 2 8 -2 9 M arch 2 0 1 2 ... 62
Table 3.7. R ayleigh lidar data acquisition param eters for 4 April 201 2 ... 66
Table 3.8. Perform ance o f the N IC T and extended lidar system s...71
Table 3.9. A ssessm ent o f the extended lidar system over 2 h ...71
Table 4.1. Characteristics o f M ILs reported in the average tem perature profile... 83
Table 4.2. Characteristics o f M ILs reported in the 2 h tem perature profile w here the am plitude is largest... 89
Table 4.3. A verage M IL characteristics for all M ILs at Chatanika, A K ...89
Table 4.4. A verage M IL characteristics o f large am plitude M ILs at Chatanika, A K ... 91
Table A.1. L idar signal statistics for the 24-inch telescope m easurem ents on 18-19 February 2012 (2245-0654 L S T )...124
List of Tables
Table A.2. L idar signal statistics for the 41-inch telescope m easurem ents on 18-19
February 2012 (2245-0654 L S T )...125
Table A.3. L idar signal statistics for the 41-inch telescope m easurem ents on 2 2-23 M arch 2012 (2133-2329 L S T )... 126
Table A.4. L idar signal statistics for the 41-inch telescope m easurem ents on 2 2-23 M arch 2012 (2359-0223 L S T )... 127
Table A.5. L idar signal statistics for 2 3 -2 4 April 2008 (0007-0215 L S T )... 128
Table A .6 . L idar signal statistics for 2 3 -2 4 April 2012 (0004-0211 L S T )... 129
Table A.7. L idar signal statistics for 24 -2 5 April 2012 (0002-0210 L S T )... 130
Table A .8 . Com parison o f pow er com pensated lidar signals for April 2008 and 2012. .131 Table A.9. L idar signal statistics for the 24-inch telescope m easurem ents on 2 8 -2 9 M arch 2012 (2259-2358 L S T )... 132
Table A.10. Lidar signal statistics for the 41-inch telescope m easurem ents on 2 8 -2 9 M arch 2012 (0035-0111 L S T )... 133
Table A.11. Set by set acquisition m ethod for 4 April 20 12 ... 134
Table A.12. L idar signal statistics for the PL8020 m easurem ents on 4 April 2012 (0024 0355 L S T )...135
Table A.13. L idar signal statistics for the PL9030 m easurem ents on 4 April 2012 (0024 0355 L S T )...136
Table B.1. SABER m easurem ents and applications... 139
Table B.2. Random and system atic errors for SABER LT E T(p) ... 148
Table B.3. Random and system atic errors for SABER n o n -L T E Tk for the upper m esosphere and low er therm osphere for m idlatitude and polar su m m e r... 153
Page
Appendix A. Chapter 3 T a b les... 124
Appendix B. SABER Temperature Retrieval in LTE and non-LTE Atmospheric R e g io n s... 137
B.1. In tro d u ctio n ...137
B.2. SABER ex perim ent... 138
B.2.1. Limb scanning m e th o d ... 139
B.2.2. SABER in strum ent... 142
B.3. SABER te m p eratu res... 143
B.3.1. CO2 15 pm b a n d s ...143
B.3.2. LT E T(p) retriev al... 146
B.3.3. N on-L T E Tk retrieval ... 148
B.3.4. O bserved therm al stru c tu re ... 153
B.3.4.1. T(p) com parison w ith Rayleigh lid a r ... 154
B.3.4.2. Tem perature com parison w ith falling sphere clim atology... 157
B.4. S u m m a ry ...158 List of Appendices
Acknowledgements
I w ould like to thank my advisor Dr. R ichard Collins for the opportunity to w ork w ith the lidar group, for his excellent advice, constant guidance, and unw avering support o f my research and involvem ent in the UAF W om en’s H ockey team. I thank my com m ittee m em bers Dr. K enneth Sassen, Dr. Ruth Lieberm an, and Dr. U m a B hatt for their advice and encouragem ent during my graduate years at UAF. The lidar data presented in this thesis w ould not be possible w ithout the dedication from students previously w orking at the L idar Research Laboratory, and current students Seth Robinson, M atthew Titus, and Cam eron M artus, and for the never ending support from Dr. Collins. I thank the staff o f P oker F lat R esearch R ange for their assistance and m aintenance o f the Lidar R esearch Laboratory. I will always be grateful to the faculty, staff, and students o f the D epartm ent o f A tm ospheric Sciences and G eophysical Institute. Special thanks to the Chair o f the D epartm ent o f A tm ospheric Sciences Dr. N icole M olders for supporting m y self and the UAF W om en’s H ockey team and for her helpful com m ents on this thesis, B arbara Day for her patience and assistance through the years, and Flora G rabow ska and all the w onderful librarians o f the K eith B. M ather Library.
I thank the N ational Science Foundation and the N ational A eronautics and Space A dm inistration for their support o f this study. Thanks to the Coupling, Energetics and D ynam ics o f A tm ospheric R egions (CED A R) program for supporting my attendance o f the 2009 and 2011 C ED A R W orkshop in Santa Fe, NM . I also w ould like to thank Dr. K ohei M izutani and N ational Institute o f Inform ation and Com m unications Technology for their support o f the L idar R esearch Laboratory.
M any thanks to B rentha Thurairajah for all her help and know ledge over the years, Jeanie Talbot and K etsiri Leelasakultum for their encouragem ent and conversation, and M atthew Titus for his help and know ledge during developm ental w ork on the lidar systems and for his com m ents on this thesis. I w ould like to thank D avid H ooper for his patience, support, inspiration, and encouragem ent during my M S research, and for his helpful com m ents on this thesis.
Last, but certainly not least, I w ould like to thank my mom and dad (Janlee and K en Irving), and my sister and brother (Bonnie and G eoffrey Irving), for their constant support and love. Thanks to the U niversity o f A laska W om en’s H ockey team for several w onderful years and special thanks our Coach, Scott V ockeroth for his dedication, and my closest friend, Shawna Jusczak, for everything.
Dedication To
Chapter 1. Introduction
1.1.
The Arctic middle atmosphere
The E arth ’s atm osphere is divided into distinct spheres separated by “pauses”, the closest o f w hich is the troposphere from the ground to ~15 km. A bove the tropopause, the stratosphere extends to the stratopause at ~50 km. A bove the stratopause the m esosphere extends to ~90 km. Finally, above the m esopause the therm osphere extends upw ard to space. The heights o f these spheres vary geographically and seasonally. F or instance, the tropopause reaches a m axim um height o f ~17 km at the equator then decreases w ith latitude until reaching a m inim um o f ~10 km at the poles [e.g., W allace and H obbs, 2006]. The study o f the vertical therm al structure o f the atm osphere contributes to the understanding o f such phenom ena as the ozone layer, noctilucent clouds (N LCs), polar stratospheric clouds (PSCs), elevated stratopause events, stratospheric sudden w arm ings (SSW s), the aurora, airglow, and m esospheric inversion layers (MILs). The vertical tem perature structure from the Extended M ass Spectrom eter Incoherent Scatter (M SISE- 90) model [Hedin, 1991] during solstice conditions at 65°N is shown in Figure 1.1 w ith regions o f interesting atm ospheric phenom enon highlighted.
Figure 1.1. Tem perature o f the E arth ’s atm osphere as a function o f height from the M SISE-90 m odel for 21 June 2011 and 22 D ecem ber 2011 at (65°N, 213°E).
The tem perature in the troposphere decreases w ith increasing height at a rate o f approxim ately 6 K km -1. This therm al structure is m aintained through a balance betw een infrared heating and radiative cooling, large-scale heat transport by large-scale w eather systems, and vertical transport o f latent and sensible heat away from the surface by small- scale turbulence. A bove the tropopause, the tem perature gradient reverses as a result o f radiative heating due to the absorption o f solar ultraviolet (UV) radiation by ozone (O3). The tem perature then decreases w ith height in the m esosphere at about the same rate as the troposphere due to reduced solar heating. In the therm osphere, tem peratures increase w ith height due to the absorption o f solar U V radiation by trace am ounts o f atomic oxygen [e.g., H olton, 2004; W allace and H obbs, 2006].
W ithout dynam ic eddy m ixing, the atm osphere w ould relax to a radiative state in w hich the tem perature w ould trail the annual solar heating cycle w ith a uniform increase from the w inter pole to sum m er pole. In this state, the circulation w ould consist o f a zonal m ean zonal flow in the therm al w ind balance w ith the m eridional tem perature gradient. In such a circulation, there w ould be no m eridional or vertical circulation, and no exchange betw een the stratosphere and troposphere. In fact, the eddy-driven circulation has both m eridional and vertical com ponents that cause large departures from the atm osphere’s radiatively determ ined state [Holton, 2004]. An exam ple o f this circulation is the cold sum m er polar m esopause and w arm w inter m esopause that is clearly evident in the tem perature profiles in Figure 1.1. From w inter to sum m er the stratopause warm s by 20 K w hile the m esopause cools by 30 K.
The m iddle atm osphere extends from above the tropopause to approxim ately 100 km, w here atm ospheric constituents are well m ixed by eddy processes. The upper boundary o f the m iddle atm osphere is near the turbopause (dashed line at 100 km in Figure 1.1), above w hich the m iddle atm osphere is dom inated by m olecular diffusion. Here, the chem ical com position o f the atm osphere varies from species to species. A lthough w eather and clim ate are prim arily due to processes occurring in the troposphere, w hich contains approxim ately 85% o f the m ass and 99% o f the w ater vapor o f the E arth ’s
atmosphere, the m iddle atm osphere also plays an im portant role. The troposphere and m iddle atm osphere are linked through dynam ical and radiative processes, and through the exchange o f trace substances im portant to the photochem istry o f the O3 layer [e.g., H olton et al., 1995].
The circulation o f the m iddle atm osphere is driven by seasonal variation in solar heating and w ave forcing, w ith w esterly zonal m ean w inds in the w inter hem isphere and easterly zonal m ean w inds in the sum m er hem isphere. A fter the autumnal equinox, the stratosphere at high latitudes cools due to the em ission o f therm al radiation and reduced incom ing solar radiation. The m eridional therm al gradient drives a pressure gradient betw een the pole and m idlatitudes that, along w ith the E arth ’s rotation, creates a circum polar belt o f w esterly w inds referred to as the polar vortex, or polar night jet. The polar vortex is a synoptic scale cyclone that isolates cold air inside the vortex and inhibits m ixing w ith w arm er m id-latitude air masses. W ithin the vortex tem peratures can fall below 195 K, allow ing PSCs to form. PSC s act as catalysts for O3 depletion by providing a surface for heterogeneous reactions to take place [Solomon et al., 1986]. Isolation w ithin the vortex further enhances O3 depletion by denitrifying the air w ithin the vortex, a necessary chem ical condition for O3 depletion [e.g., Toon and Turco, 1991; Schoeberl and H artm ann, 1991].
W ith the discovery o f the A ntarctic ozone hole in the 1980’s [Farman et al., 1985], there w as a surge o f scientific interest and research regarding O3 depletion and m iddle atm ospheric circulation. O3 is the prim ary reason the E arth’s surface is not constantly bom barded w ith U V radiation from the sun. The A ntarctic O3 hole is rem otely located, but as the vortex splits and dissipates in the spring the depleted air is carried to low er latitudes and poses a health threat [e.g., A tkinson et al., 1989]. A pproxim ately 70% o f the O3 above the A ntarctic (approxim ately 3% o f the E arth ’s total) is lost during Septem ber and O ctober [Toon and Turco, 1991]. In spring 2011, unprecedented ozone depletion that reached levels sim ilar to the A ntarctic ozone hole w as recorded in the A rctic [M anney et al., 2011]. The record ozone loss has attracted considerable attention
due to its im plications for a m ajor health risk in the highly populated N orthern H emisphere.
The A rctic m iddle atm osphere has gained w idespread attention in recent years due to its rem arkable inter-annual variability [e.g., M anney et al., 2005; 2009; 2011]. A lthough initiatives focused on the A rctic such as the International Polar Y ear (IPY) [Collins, 2004; ICSU, 2004; N RC, 2004] have strengthened our know ledge o f the atm ospheric circulation in this region, our present understanding o f w hat drives such variability is incom plete [Newm an et al., 2001; M anney et al., 2011]. The study o f phenom ena such as SSWs, elevated stratopause events, and M ILs are necessary to developing a thorough understanding o f the m iddle atm ospheric [e.g., Fritts and A lexander, 2003; M eriw ether and Gerrard, 2004; Chandran et al., 2011; de la Torre, 2012].
1.2.
The wave-driven circulation of the middle atmosphere
A tm ospheric w aves play a m ajor role in the circulation o f the m iddle atm osphere and have a profound influence on the tem perature structure. W aves can propagate vertically and horizontally from w ave sources to regions w here transience, nonlinear w ave breaking, or dissipation causes m om entum transfer to the m ean flow [Holton and Alexander, 2000]. W ave transience is the local grow th or decay o f w ave am plitude, and dissipation can be either radiative or turbulent. A tm ospheric w ave m otions result from a balance betw een inertia and restoring forces acting on fluid parcels displaced from their equilibrium longitudes or altitudes [Holton and A lexander, 2000]. Gravity w aves, som etim es referred to as buoyancy w aves, are small scale w aves w ith horizontal w avelengths on the order o f ten to hundreds o f kilom eters w hose restoring force is buoyancy. Topography is one o f the m ain sources o f vertically propagating gravity w aves, know n as m ountain or topographic waves. Planetary-scale Rossby w aves (henceforth referred to as planetary w aves) are generated by large-scale orography and land-sea contrast. They are large scale w aves, on the order o f ~104 km, w hose restoring force is the variation o f the Coriolis param eter w ith latitude, or the m eridional gradient o f
potential vorticity know n as the P-effect [e.g., Salby, 1996; H olton and A lexander, 2000; H olton, 2004].
The phase speed o f planetary w aves is w estw ard relative to the m ean flow, and can only propagate vertically in the w inter hem isphere w hen the w esterly w inds are w eaker than a critical value (e.g., u = 0 for quasi stationary planetary waves), dependent on the horizontal scale o f the w aves [Charney and D razin, 1961]. F or planetary w aves, w ave breaking represents an irreversible process in w hich the w ave deposits its m om entum and energy into the system. U pw ard propagating planetary w aves grow in am plitude as the atm ospheric density decays w ith height. Therefore, at some altitude the disturbance am plitude will becom e so large that overturning and dissipation m ust occur. The irreversible deform ation o f m aterial contours during planetary w ave breaking is responsible for the stratospheric surf zone (dashed lines in Figure 1.2) [M cIntyre and Palm er, 1984], has been shown to cause SSW s [M atsuno, 1971], and is one o f the proposed form ation m echanism s o f M ILs [Wu, 2000; Salby et al., 2002].
The zonal phase progression o f gravity w aves can be either w estw ard or eastw ard w ith parcel oscillations perpendicular to the phase progression. A nalogous to planetary waves, gravity w ave am plitudes strengthen in the m esosphere and instability can occur w hen the am plitude disturbance becom es large. W ave breakdow n caused by convective instability can m ix through deep layers and may lead to substantial changes in the local chemical com position [Holton and A lexander, 2000; Fritts and A lexander, 2003]. Gravity w ave breaking forces accelerations in the background flow and “drags” the flow tow ard the phase speed o f the wave. The filtering o f vertically propagating gravity w aves by stratospheric w inds allows w estw ard propagating w aves to propagate through the w inter stratosphere, w hile only eastw ard w inds w ith zonal phase speeds larger than the m axim um w ind speed are transm itted. The opposite is true for the sum m er stratosphere, and results in an eastw ard drag force exerted in the sum m er m esosphere and a w estw ard drag force in the w inter m esosphere. In the m esosphere, the m ean zonal w ind distribution
is m aintained by the m eridional drift, a consequence o f this strong gravity w ave forcing [Holton and Alexander, 2000], shown in Figure 1.3.
L a titu d e °N
Figure 1.2. Schem atic illustration o f January-February stratospheric circulation and eddy heat fluxes (adapted from N ew m an et al. [2001]). Shown are regions o f (a) planetary w aves propagating into the stratosphere, (b) slowly refracting tow ard the equator, (c) depositing easterly m om entum , and (d) inducing a residual circulation that causes uplift in the tropics and sinking in the polar region. Short arrows illustrate the w ave propagation and the thick line w ith arrow s shows the residual circulation. The thin solid lines show the w ind speed, the dotted line shows the tropopause, and the dashed line shows the EP flux divergence.
S u m m er W in te r 80 ■g 70 3 < 60 -9 0 -6 0 -3 0 0 30 60 90 L atitu d e
Figure 1.3. Schem atic illustration o f the solstice season pole-to-pole w ave-driven transport circulation in the m esosphere (adapted from H olton and A lexander [2000]). Shading shows regions o f zonal forcing by gravity w ave breaking.
Planetary w ave forcing leads to an equator-to-pole dow nw ard circulation in the w inter hem isphere (Figure 1.2), know n as the B rew er-D obson circulation. G ravity w ave forcing leads to a pole-to-pole circulation (Figure 1.3) from the sum m er to w inter pole [Houghton, 1978]. In the m iddle atmosphere, the planetary w ave forced circulation is responsible for transport o f critical species such as O 3 and w ater vapor, but is absent in sum m er due to easterly m ean winds. The gravity w ave induced circulation is responsible for the cooling o f the polar sum m er m esopause and the w arm ing o f the polar w inter stratopause, visible in Figure 1.1 [e.g., H olton and A lexander, 2000; Fritts and A lexander, 2003]. M axim um O3 loss in the A rctic polar vortex coincides w ith the return o f sunlight in M arch, w hich initiates the photolysis o f O3. The interannual variability o f M arch stratospheric tem peratures in the A rctic is closely related to the January-February eddy heat flux caused by planetary w aves [Newm an et al., 2001]. The eddy heat flux is also correlated w ith the stratospheric m om entum flux, Eliassen-Palm er (EP) flux divergence, and zonal m ean wind. E P flux is a representation o f the atm ospheric eddy-m ean flow
interaction and provides a m easure o f the zonal-forcing o f the zonal m ean flow. N egative E P flux divergence corresponds to the w estw ard zonal force exerted by eddies on the atm osphere [Holton, 2004].
Figure 1.2 is a schem atic illustration o f the N orthern H em isphere January-February eddy heat flux adapted from N ew m an et al. [2001]. U pw ard propagation o f planetary w ave activity into the stratosphere (Figure 1.2a) is represented by large eddy heat fluxes. As planetary w aves propagate upw ard, they encounter the polar vortex and can be refracted equatorw ard w here they deposit easterly m om entum (Figure 1.2b). This m om entum deposition is balanced by a northw ard residual circulation (Figure 1.2c) w hich decelerates the polar night jet. This w ave induced m eridional residual circulation causes rising m otion in the tropics and sinking m otion at the poles (Figure 1.2d) [see N ew m an et al., 2001 and references therein]. Thus, if m idw inter planetary w ave activity is w eak, there is little deceleration o f the polar night jet, w hich leads to a cold stable polar vortex in M arch. A stronger and m ore stable vortex in the spring enables tem peratures to fall low enough for m ore PSCs to form and for longer periods o f time, enhancing O3 depletion [e.g., Toon and Turco, 1991; Schoeberl and Hartm ann, 1991].
The A rctic m iddle atm osphere is difficult to characterize due to the nonlinear w ave-w ave and w ave-m ean-flow interactions that result in the zonal asym m etric w intertim e circulation. SSWs, an exam ple o f the w ave-m ean flow interaction, are caused by the interaction betw een upw ard propagating planetary w aves and the zonal polar stratospheric flow. The m echanism for SSWs w as first proposed by M atsuno [1971] and their occurrence is characterized by a displacem ent o f the polar vortex, w eakening o f the zonal m ean zonal flow, and an asym m etric stratospheric circulation. SSWs are defined as m ajor w hen at the 10 hPa level, or below, the zonal m ean tem perature increases polew ard o f 60°N and the zonal m ean zonal w ind reverses [Labitzke, 1972] and m inor w hen no w ind reversal takes place. A contem porary exam ple o f an A rctic SSW was presented by Chandran et al. [2011] using a case study approach o f a m ajor SSW (Figure
1.4) follow ed by an elevated stratopause in the free running W hole A tm osphere Com m unity Clim ate M odel (W ACCM ).
Q uiet M odel Year - 1996/1997 D isturbed M odel Year - 1973/1974
(a)
Dec 11 Dec 31 Jan 20 Feb 9 M a r l Dec 11 D ec 31 Jan 20 Feb 9 M a r l
D ec 11 Dec 31 Jan 20 Feb 9 M a r l D ec 11 Dec 31 Jan 20 Feb 9 M a r l (C)
Dec 11 Dec 31 Jan 2 0 Feb 9 M a r l Dec 11 D ec 31 Jan 20 Feb 9 M a r l (d)
Dec 11 D ec 31 Jan 20 Feb 9 M ar 1 Dec 11 D ec 3 1 Jan 20 Feb 9 M ar 1
Day o f the Year Day o f the Year
Figure 1.4. The zonal m ean (a) tem peratures at 70°N w ith dashed line indicating start o f a SSW, (b) wind, (c) EP flux divergence, and (d) gravity w ave forcing averaged for 5 5 - 70°N for (left) model year 1996/1997 and (right) model year 1973/1974 plotted as a function o f altitude and day. In the bottom three panels, red (positive) contours denote eastw ard flow /forcing and blue (negative) denote w estw ard (adapted from Chandran et al. [2011]).
D uring the dynam ically quiet w inter in the W A C C M model year 1996/1997 (left) the zonal m ean tem perature (Figure 1.4a) behaves as a standard atmosphere, w ith a m axim um tem perature at the stratopause near 55 km and tem peratures decreasing through the m esosphere. The zonal m ean w ind (Figure 1.4b, left) shows a strong eastw ard je t in the stratosphere w ith speeds greater than 60 m/s and w inds reversing near
80 km to w estw ard w inds in the m esosphere. There is w eak planetary w ave forcing (Figure 1.4c, left) throughout the year w ith m id-w inter EP flux divergence values generally less than ±10 m/s/d. W estw ard gravity w ave forcing (Figure 1.4d, left) reaches a m axim um o f ~50 m /s/d in the m esosphere betw een 60 and 80 km. The W A C C M model year 1996/1997 represents a dynam ically quiet year w here the stratospheric je t rem ains strong and eastw ard w ithout reversals, contrary to 1973/1974 (Figure 1.4b, right).
The zonal m ean tem perature for the disturbed W A C C M model year 1973/1974 (Figure 1.4a, right) shows a stratopause near 60 km through the beginning o f Decem ber. On D ecem ber 11th the stratopause w arm s and low ers in altitude to 46 km and on D ecem ber 26th the stratopause again w arm s w ith m axim um tem peratures at 45 km. B oth w arm ing events are follow ed by a cooling o f the stratopause and then a nearly isotherm al atm osphere betw een 30 and 75 km from approxim ately January 10-January 17. Preceding both SSW is strong w estw ard planetary w ave forcing, w ith EP flux divergence values reaching -60 m/s/d. The strong w estw ard planetary w ave forcing reverses the stratospheric eastw ard je t to w estw ard and induces a polew ard and dow nw ard circulation, leading to adiabatic w arming. The w estw ard stratospheric je t allows for eastward propagating gravity w aves to penetrate through the stratosphere and the eastw ard gravity w ave forcing (Figure 1.4c, right) then reverses the m esospheric w inds from w estw ard to eastward. The eastw ard w inds in the m esosphere result in an equatorw ard and upw ard flow. An elevated stratopause forms near 70 km follow ing the SSW s and gradually relaxes down to ~55 km by M arch 1st. A s w estw ard gravity w ave forcing returns and strengthens after -Jan u ary 17th, polew ard and dow nw ard flow is re-established and the stratopause w arm s and low ers through the w ave-drive diabatic descent.
The study by Chandran and co-w orkers dem onstrates the com plexity and variability that w ave-w ave and w ave-m ean flow interactions, and their nonlinear feedbacks have on the A rctic m iddle atm osphere [Chandran et al., 2011]. D irect observations o f the w ave- driven m iddle atm osphere are essential in furthering our understanding o f the
dynam ically driven Arctic, and to the developm ent o f m odels such as W ACCM . Located in the W estern Arctic, Chatanika is a natural laboratory w here the understanding o f w ave- breaking, turbulence, and instability is advanced. Its location provides a unique opportunity to study w ave-m ean flow interactions, interactions betw een large and small- scale dynam ical processes, and feedbacks from sm all-scale processes on the large-scale circulation in a region w here the w intertim e circulation is greatly influenced by the A leutian anticyclone and polar vortex. R ecent studies have found that during periods o f strong planetary w ave am plitudes and negative EP flux divergence, interactions betw een the A leutian anticyclone and the polar vortex lead to their irreversible intertw ining [Harvey et al., 2002; Thurairajah et al., 2010a; 2010b; Chandran et al., 2011]. These findings point to the W estern A rctic atm osphere as a hub for planetary w ave activity and subsequent planetary w ave breaking.
M ILs are understood to be a signature o f nonlinear w ave-w ave and w ave-m ean-flow interactions and are im portant to study for tw o prim ary reasons: stability and energy transfer. The m echanism responsible for the form ation o f M ILs is still am biguous, although many have been proposed [see recent review by M eriw ether and Gerrard, 2004]. M ILs are identified as a layer o f increasing tem perature in the m esosphere, w here tem peratures typically decrease w ith height [Schmidlin, 1976]. M ILs are so nam ed because their appearance is sim ilar to inversions capping the diurnally driven atm ospheric boundary layer [M eriw ether and Gerrard, 2004]. As tem peratures increase on the bottom side o f the M IL, atm ospheric stability is increased and vertical m ixing is reduced. As tem peratures decrease on the topside o f the M IL, atm ospheric stability decreases and can approach instability at the adiabatic lapse rate. The instability above a M IL can lead to convective and/or dynam ic instability and support the developm ent o f turbulence. Indeed, M ILs w ith topside lapse rates (-3T /3 z) approaching the adiabatic lapse rate (-9 .8 K km -1) have been reported [e.g., W hitew ay et al., 1995; Cutler et al., 2001; D uck et al., 2001; D uck and Greene, 2004]. The generation o f turbulence in the w intertim e atm osphere is im portant in quantifying the im pact o f space radiation on the atmosphere. Current model studies show that turbulent transport is necessary, in addition
to the dow nw ard w ave-driven circulation, for the transport o f nitrogen com pounds produced by energetic particle precipitation from the therm osphere into the m esosphere w here they contribute to ozone depletion [Smith et al., 2011]. This transport is critical in understanding how m eteorological processes control the im pact o f space processes on the E arth ’s atmosphere.
U nderstanding energy transport in the m iddle atm osphere provides further m otivation for the investigation o f M ILs. In addition, M ILs have a profound im pact on the propagation o f gravity w aves in the m esosphere [e.g., Taylor et al., 1995; D ew an and Picard, 1998; D ew an and Picard, 2001]. A tm ospheric w ave breaking is the prim ary m ethod by w hich energy is transferred up from the low er atm osphere and has been proposed as the form ation m echanism o f M ILs [e.g., M eriw ether and Gerrard, 2004; Gan et al., 2012 and references therein]. Investigation o f M ILs has historically relied on satellite and lidar data and a com prehensive understanding o f the phenom ena is incom plete. M ILs have been reported from the equator and subtropics [e.g., D ao et al., 1995; Fechine et al., 2007] to high latitudes [e.g., C utler et al., 2001; D uck and Greene, 2004], and w ith lifetim es ranging from days [e.g., Salby et al., 2002; Gan et al., 2012] to hours [e.g., Collins et al., 2011]. The geographic and tem poral variability o f M IL characteristics point to m ultiple form ation m echanism s and illustrates the difficulty in elucidating their com plete physical understanding.
1.3.
Rayleigh lidar
L idar is an acronym for LIight R anging A nd D etection [M iddleton and Spilhaus, 1953]. A nalogous to optical radar, lidar is a form o f active rem ote sensing that m easures the laser light backscattered from a target m olecule or atom. R ayleigh lidar systems m easure the Rayleigh scattered light from air m olecules in the cloud and aerosol free region o f the m iddle atmosphere. R ayleigh scatter is defined as the scattering o f electrom agnetic radiation by particles sm aller than the w avelength o f radiation [Strutt, 1899].
B efore the technique o f tem perature m easurem ent by R ayleigh lidars w as developed, the - 3 0 to 90 km altitude range accessible to R ayleigh lidar posed a challenge to researchers interested in the aeronom y o f the m iddle atm osphere [H auchecorne and Chanin, 1980]. W eather balloons pop at -2 5 km, nadir pointing satellites are generally lim ited to below 55 -6 0 km and above 200 km, radars face a lack o f scattering media, airglow im agers observe only thin layers (e.g., the hydroxyl (OH) layer from - 8 4 to 94 km [Baker and Stair, 1988; B rinksm a et al., 1998]), and rockets are expensive and provided only instantaneous m easurem ents. In 1938, cloud base heights w ere determ ined for the first tim e using pulses o f light [Bureau, 1946]. Elterm an [1951] used a searchlight to sound the low er and m iddle atm osphere, and w as the first to use the lidar integration technique to m easure atm ospheric tem peratures. K ent et al. [1967] reported the first m easurem ents o f atm ospheric constituents m ade w ith a Q -sw itched laser, and H auchecorne and Chanin [1980] w ere the first to obtain a tem perature m easurem ent betw een 35 and 70 km. Today, R ayleigh lidar is a robust m easurem ent technique and has contributed extensively to the understanding o f the m iddle atm ospheric com position and circulation [W eitkamp, 2005]. F urther inform ation on lasers is available in V erdeyen [1981] and Silfvast [1996], and on lidar m ethods and applications in atm ospheric science in M easures [1984], Fujii and Fukuchi [2005], and W eitkam p [2005].
The N ational Institute o f Inform ation and Com m unications Technology (N ICT) Rayleigh lidar observations at P oker Flat R esearch Range, Chatanika, A laska (65°N, 213°E), have yielded density and tem perature m easurem ents from 4 0 -8 0 km. These m easurem ents are distributed betw een A ugust and M ay and have been m ade under clear nighttim e skies since N ovem ber 1997. R ayleigh lidar observations at P FR R are used to study the tem perature structure over Chatanika [e.g., Thurairajah et al., 2009], as well as the w ave- driven circulation and turbulence [e.g., W ang, 2003; Collins and Smith, 2004; Thurairajah et al., 2010a; 2010b; Collins et al., 2011; N ielsen et al., 2012], and phenom ena such as N LC s [e.g., Collins et al., 2003; Collins et al., 2009; Taylor et al., 2009; K elley et al., 2010; V arney et al., 2011].
1.4.
Scope of this study
In this thesis I present a scientific study o f M ILs observed by lidar and a technical study o f the extension o f the lidar system at Chatanika. Based on these studies, I show how the extended lidar system w ill broaden the scope o f future studies at Chatanika.
In C hapter 2, I review the principles and technique o f Rayleigh lidar. I present the lidar equation used for tem perature retrievals from 40 km to 80 km and its application to study atm ospheric structure over Chatanika. The N IC T Rayleigh lidar system and system perform ances at Chatanika are reviewed.
In Chapter 3, I present results o f the field tests I conducted during the extension o f the N IC T R ayleigh lidar system in spring 2012. I assess the ability o f this new lidar system, w hich incorporates a larger aperture telescope and higher pow er laser, to extend the scope o f future studies at Chatanika. I investigate the nonlinear response in the detector electronics below 50 km and their effects on the lidar tem perature retrieval.
In Chapter 4, I present a clim atology o f M ILs observed at C hatanika based on observations m ade over a tw elve year period. Their relationship to the planetary w ave structure is exam ined based on satellite m easurem ents m ade over a six year period. G eopotential height and tem perature data from the Sounding o f the A tm osphere using B roadband Em ission R adiom etry (SABER) instrum ent on the Therm osphere Ionosphere M esosphere Energetics D ynam ics (TIM ED ) satellite are used to study the planetary w ave field at 65°N.
In Chapter 5, I sum m arize the key findings o f my study and present conclusions. D irection and suggestions are given for future studies o f M ILs and developm ent o f the R ayleigh lidar at the Lidar R esearch Laboratory, PFRR, Chatanika, Alaska.
Chapter 2. Principles and Techniques of Rayleigh lidar
2.1.
NICT Rayleigh lidar
The N ational Institute o f Inform ation and Com m unications Technology (N ICT) Rayleigh lidar w as deployed at Poker F lat R esearch R ange (PFRR), Chatanika, A laska (65°N, 213°E) in N ovem ber 1997. The N IC T R ayleigh lidar w as deployed by researchers from N IC T and the G eophysical Institute o f the U niversity o f A laska Fairbanks (G I-U A F ) as part o f the A laska project. The goal o f the A laska project w as to establish a com prehensive m easurem ent system capable o f enduring and docum enting the A rctic environm ent [M izutani et al., 2000; M urayam a et al., 2007]. The R ayleigh lidar system w as initially installed in the Optics Laboratory (OL) at PFRR. The R ayleigh lidar w as relocated to the D avis Science Center, PFRR, in April 1999 w hile the OL w as dem olished and the L idar R esearch Laboratory (LRL) w as built on the OL site. The lidar w as installed in LRL in July 2000 and has operated there since.
A detailed schem atic o f the N IC T lidar system is shown in Figure 2.1 and the specifications o f the system are listed in Table 2.1. The N IC T R ayleigh lidar transm itter consists o f a laser, a laser beam expander (BE), a beam steering m irror (BSM ), and a laser pulse detector (LPD), consisting o f a laser diode. The N IC T lidar system laser is a N d:Y A G Continuum ® Pow erlite 8020 Q -sw itched laser operating at a repetition rate o f 20 pulses-per-second (pps) and w ith a w avelength o f 532 nm. F urther inform ation on N d:Y A G lasers can be found in current textbooks [e.g., V erdeyen, 1981; Silfvast, 1996]. The telescope used in the receiver system is a classic N ew tonian telescope w ith a 62 cm diam eter (further inform ation on astronom ical optics can be found in Schroeder [2000]).
Table 2.1. N IC T Rayleigh lidar system specifications. Transm itter
L aser N d:Y A G
M odel Continuum Pow erlite 8020
W avelength (1L) 532 nm R epetition R ate (RL) 20 Hz
Pulse Energy (EL) 375 - 460 mJ
Pulse W idth 5 - 7 ns Line W idth 1.0 cm -1 (28 pm, 30 GHz) Beam Expander x 10 D ivergence 0.45 m rad Receiver Telescope N ew tonian
O uter D iam eter 620 mm
Inner D iam eter 200 mm
Collecting A rea 0.270 m2
Range Resolution 75 m
Optical B andw idth 0.3 nm Field o f view (FOV) 1 mrad
D etector Photom ultiplier Tube
M odel H am am atsu R 3234-01
Pulse duration 5 ns
D ark Count 50 - 150 counts/second Pream plifier Gain x 5
M odel Stanford R esearch Systems SR445
Bandw idth DC to 300 M Hz
D igital R ecorder M ultichannel Scalar
M odel Ortec Turbo M CS T914
Figure 2.1. Schem atic diagram o f the N IC T R ayleigh lidar system at PFRR, Chatanika, A laska (65°N, 213°E).
BC - Blanking Control BE - Beam Expander
BSM - Beam Steering Mirror CL - Collimating Lens
HVPS - High Voltage Power Supply IF - Interference Filter
LPD - Laser Pulse Detector
MCS - MultiChannel Scalar PA - Pre-Amplifier
PH - Pin Hole PM - Primary Mirror
PMT - Photo Multiplier Tube SM - Secondary Mirror
The N d:Y A G laser em its a pulse o f light in a beam that passes through the beam expander (BE). The expanded beam is reflected by the beam steering m irror (BSM ) into the sky. Photons o f light that are Rayleigh scattered in the atm osphere are detected by the telescope. The light incident on the telescopes prim ary m irror (PM ) is reflected to the secondary m irror (SM). The light is reflected from the SM through the pinhole at the focal point o f the telescope. The use o f the pinhole defines the telescopes field-of-view (FOV). The light w ithin the telescopes FO V passes the pinhole and is collim ated by the collim ating lens (CM). The telescope FO V is determ ined by the pinhole (PH) placed in front o f the photom ultiplier tube (PMT). The collim ated light at 532 nm passes through the interference filter (IF) w here it is focused on the PM T. Note: previous system descriptions [e.g., W ang, 2003; N adakuditi, 2005] included a beam splitter before the
interference filter. The beam splitter w as taken out in early 2008 w hen Resonance lidar m easurem ents w ere m ade through a separate telescope [Light, 2009].
The transm itter com m unicates w ith the receiver system through the laser pulse detector and the lasers Q-switch. The Q -sw itch signal triggers the blanking control (BC) to transm it a pulse o f 150 ps duration (know n as a gate pulse) to the PM T (a general discussion o f PM Ts can be found in H am am atsu Photonics [2005; 2006]). The gate pulse biases the first dynode o f the PM T by +200 V (H VPS2) above its normal operating voltage o f -2000 V (HVPS1), reducing the gain o f the PM T by a factor o f greater than 106. The gate pulse acts as electronic blanking causing return signals from altitudes below ~25 km to be detected at low gain and allows for the detection o f high altitude signal returns at high gain. The LPD detects the transm itted laser pulse and triggers the m ultichannel scalar (M CS), a high-speed counter, to count the incom ing PM T pulses in a given tim e w indow at rates up to 150 MHz. The tim e w indow determ ines the spatial resolution. F or exam ple, a bin tim e o f 0.5 ps results in a 75 m vertical resolution. The current signal from the PM T is first am plified by the pre-am plifier (PA) before being sent to the M CS. The M CS typically records 4096 range bins, resulting in an echo profile from the ground to ~300 km. The next laser pulse triggers the next M CS profile acquisition w hich is added coherently to the previous one. The profiles are added together for a predeterm ined num ber o f laser pulses, typically 1000 or 2000, to yield a single raw data profile. The single raw data profile is transferred to the com puter w here it is stored and then initiates the M CS to begin a new profile on the next laser pulse.
N ighttim e lidar observations at PFRR are taken under clear sky conditions betw een A ugust and M ay w ith the m ajority o f observations taken betw een O ctober and M arch. These observations have yielded tem perature m easurem ents o f the A rctic middle atm osphere [Thurairajah et al., 2009]. N o tem perature m easurem ents are taken from mid M ay through late July due to the background light levels in sum m er tw ilight at Chatanika. L idar retrieval m ethods have been developed by students at UAF to allow m easurem ents at different resolution to yield robust characterizations and estim ates o f the
geophysical variability o f the A rctic m iddle atm osphere [e.g., Cutler, 2000; W ang, 2003; N adakuditi, 2005; Thurairajah, 2009].
2.2.
The lidar equation
The lidar signal is com posed o f a sequence o f pulses that is governed by photon counting statistics. The expected returned lidar signal is proportional to the atm ospheric density. U nder clear sky conditions, the expected total lidar signal from an altitude range (z-Az/2, z+Az/2) in a tim e interval At is given by the lidar equation,
NTqT{z~) = ^s(.z ~) + n b + (2 1)
w here NS(z) is the lidar signal count proportional to the atm ospheric density, NB is the background skylight count, and ND is the detector dark count given by,
Nb = ^ HNRLA t n [ — ^ ) A t AA 2Az the ~ / ~ A
(2.3)
( 2Az \
ND = (CNRLA t ) { — ) (2 .4)
In the above equations, n is the receiver efficiency, XL is the transm ission w avelength (m), T is the atm ospheric transm ission at XL, E L is the laser energy per pulse (J), R L is the laser repetition rate (s-1), p(z) is the num ber concentration o f scatterers at an altitude z (m" 3), A t is the telescope area (m2), a R is the effective backscatter cross section at XL (5.22*10-31 m 2), h is P lanck’s constant (~6.63*10-34 J s), c is the speed o f light (~3.0*108 m s-1), A 0 r is the FO V o f the receiver (radians), HN is the background sky irradiance (W m -2 pm sr), AX is the bandw idth o f the detector (pm), and CN is the dark count rate for the detector (s-1). Standard atm osphere densities at 20 km, 40 km, 60 km, and 80 km are
8 .8 x 10-2, 3.9x10"3, 2.9x10"4, and 1.6 x 10'5 kg m-3 respectively. Pressures at these altitudes are 5.5X101, 2.8x10°, 2 .0 x 10-1, and 8 .9 x 10-3 hPa, respectively [USSA, 1976].
Figure 2.2 shows the total lidar signal profile, Ntot(z), as a function o f altitude m easured by the N IC T R ayleigh lidar system on 2 -3 January 2012. The laser pulse energy w as 400 mJ. The total lidar signal represents the signal collected from 784,000 laser pulses over an 11.7 h w indow from 2127-0913 LST (LST = U T - 9 h) on the night o f 2 -3 January 2012. B etw een approxim ately 25 km and 90 km the profile decays w ith altitude as the density o f the atm osphere decreases w ith altitude. A bove ~90 km, the signal is dom inated by background skylight, N b, and detector dark signal, N d . B elow 25 km the lidar signal is reduced due to electronic switching o f the lidar receiver detector to avoid the high signal returns from the dense low er atmosphere.
Figure 2.2. Photon count profile plotted as a function o f altitude m easured by the N ICT R ayleigh lidar at P FR R on the night o f 2 -3 January 2012. The profile w as acquired by integrating the signal from 784,000 laser pulses over a period o f 11.7 h.
W hile Figure 2.2 shows the lidar signal integrated over the w hole observation period, the raw lidar signals are acquired in the follow ing fashion. As described above, the lidar signal from several (typically 1000) laser pulses is integrated over a short period
(typically 50 s) period in the M CS unit and stored as a single raw data photon count profile. The profile is then transferred to the com puter and then displayed on the com puter screen. The com puter initiates a new profile. This process continues for several profiles (typically 16) called a set. Each set is associated w ith an individual data file. The data files are nam ed by date and sequentially by set num ber (e.g., the first set on January 2 is nam es JA02RY .001). Once a set is com pleted the data acquisition program pauses for the operator to review the data, m ake system adjustm ents (e.g., adjust laser energy, adjust interference filter, pause acquisition due to increasing cloud cover) as required, and begin the next set. Each set takes less than 14 m inutes to com plete and typically 50 sets are acquired over the course o f a 12 h observation period. On 2 -3 January 2012 the observation period lasted 11.7 h w hile the laser operated for 10.9 h, w ith 0.8 h o f tim e spent on m aking adjustm ents to the lidar system during the observation.
U nder clear sky conditions, the lidar signals rem ain relatively constant through the observation and provide uniform m easurem ents o f density and tem perature through the w hole observation period. The operator uses the lidar signal per-laser-pulse as an operational indicator o f signal quality. The lidar signal per-laser-pulse is calculated as the lidar signal profile over a certain num ber o f laser pulses divided by the num ber o f laser pulses (e.g., 1000 for a single raw data profile or 16000 for a data set). U nder good conditions (i.e. clear skies and m axim um laser energy) the lidar signal per-laser-pulse o f one photon count per-laser-pulse is expected from the 60-65 km altitude region. Signals low er than this indicates reduced atm ospheric transm ission due to clouds and/or aerosols, laser pulse energy, or receiver efficiency. Figure 2.3 shows the lidar signal per-laser- pulse calculated for each set on the night o f 2 -3 January 2012. The average total signal per-laser-pulse from 60 -65 km w as 0.84 on 2 -3 January 2012. W hile the signals rem ain constant over the first 45 sets (2126-0816 LST) the signals at all altitudes rise steadily over sets 46 through 49 (0817-0914 LST). The rise is m ost pronounced in the signal at the higher altitudes (80-85 km, 9 0-95 km, and 100-105 km ) after set 45. This increase in signal begins at an approxim ate solar angle o f -12° [USNO, 2012] and reflects the fact
that the total signal is dom inated by the background signal at these higher altitudes. The data also shows the lim itations o f using the N IC T R ayleigh lidar system for daytime observing. 2-3 January 2012 0) 1 0 3 _v> 3 L. 1 0 Z ts>
s
N 4 - 1 0 1 E O E O 1 0 ° 4— iE ^ - 1 1 0 E <31 -i o 1 0 o 1 - i o 3 -III 1 1 1 1 1 1 i i i 1 1 1 1 1 1 1 1 1 , , , 1 1 1 1 1 1 I I 40-45 km : 60-6 jSQOC 5 km g-BHB*JQ13B1 reoot : 30-35 km i i i t l !*»*■ 90-95 km 100-105 km ... . . . i . . . i . . 7 i . . . i . . . ... . i 9 13 17 21 25 29 33 37 41 45 49 Set NumberFigure 2.3. Signal levels per-laser-pulse as a function o f set num ber on 2 -3 January 2012 (2127-0913 LST).
The com bined background skylight and dark signal (Nb + N D) is 44 photon counts on 2-3 January 2012. This signal is evident as the value o f the lidar signal that rem ains constant w ith altitude above 110 km in Figure 2.2. The specific value is calculated as the average o f the lidar signal over 5 km altitude range centered around 225 km. The PM T has typical anode dark counts for the PM T o f 50 s-1 [H am am atsu Photonics, 2006]. In a 75 m range bin, this w ould result in a signal o f 2 .5 M 0 "5 photon counts per-laser-pulse. Thus, the dark current contributes a dark signal of 20 photon counts to the total lidar signal (Figure 2.2) and a signal o f 1.7x10"3 to the total signal per-laser-pulse (Figure 2.3). D uring nighttim e conditions (corresponding to astronom ical darkness w ith the sun tw elve or m ore degrees below the horizon), the background skylight and dark detector signals have sim ilar contributions to the total lidar signal.
The photon counting process acts as a Poisson random variable [e.g., Papoulis, 1984; Taylor, 1996]. F or Poisson random variables, the variance equals the expected value and
1/2
there is an inherent standard deviation Ntot for an expected photon count signal o f value Ntot. The relative error in the total signal photon count is given by,
ANTOT TOT
NTOT N
(2.5)
TOT VNTOT
and the signal photon counting error is,
ANS ANTOT VN t o t + Nb + Nc
NS N c Nc
(2 .6)
In an ideal lidar w orld (i.e. a perfect detector and no skylight) Nb and Nd w ould be zero and from Equation 2.6 w e can see that the relative uncertainty in the photon counting
1/2
signal decreases to ANs/Ns = Ns" . The lidar relative error profile for 2 -3 January 2012, expressed as a percentage, is shown in Figure 2.4. The relative error increases w ith altitude as the lidar signal reduces in height.
Figure 2.4. R elative photon counting error (%) plotted as a function o f altitude from 2 -3 January 2012.
The relative errors in the photon count profile in Figure 2.2 are calculated for 13 altitudes in Table 2.2. The signals are averaged over 5 km altitude and thus represent the statistically robust signal in a 75 m range bin at the center o f the altitude range. The
relative error increases from less than 0.2% at 40 km to over 100% at 100 km. The error increases m ore rapidly w ith height as the total lidar signal is dom inated by the com bined background skylight and dark detector signals. A bove 92.5 km, the com bined background and dark signals are larger than the atm ospheric density com ponent o f the lidar signal and the relative errors are greater than 30%. sim ila r analyses can be found in W ang [2003] and N adakuditi [2005] for 2 0-21 D ecem ber 2002 and 7 -8 M arch 2002, respectively.
Table 2.2. Lidar signal statistics for 2 -3 January 2012 (2127-0913 LST).
Altitude (km) Total Signal, Nt o t 1 ,2 (Photon count) Signal, Ns 1 ,2 ,3 ,4 (Photon count) Relative Error ANs/Ns (%) 42.5 296658 296614 1.8X10"1 47.5 122868 122824 2.9x10"' 52.5 53004 52960 4.3x10"' 57.5 23775 23731 6.5x 10"1 62.5 9899 9855 1.0 x 10° 67.5 3933 3889 1.6 x 10° 72.5 1505 1461 2.7x 100 77.5 529 485 4.7x 100 82.5 213 169 8.6 x 100 87.5 110 66 1.6 x 101 92.5 72 28 3.0x 101 97.5 57 13 5.8x 101 102.5 50 6 1.2 x 102
1: Signals are averaged over 5 km altitude and represents signal in 75 m range bins 2: Numbers are rounded to the nearest whole number
3: Background signal, Nb + Nd = 44 photon counts 4: 784,000 laser pulses transmitted
From E quation 2.2 w e can see that all com ponents o f the lidar system are directly proportional to the tem poral resolution, At, and the spatial resolution, Az, and can be rew ritten in term s o f system atic constants defined by Equations 2.10-2.12.
Ns(z) = K s ^ ( A z A t ) (2.7) Nb = KBHN(AzAt) ND = K D(AzAt) (2 .8) (2.9) W here, Ks = n T2^ aRAT (210) Kb = nRLn (A| R) AtAA (2. " ) Kd = CnRl(2 ) (212)
By integrating the photon count profiles in tim e and/or space, the product o f the spatial and tem poral resolution increases by a factor k (i.e., AtAz ^ kAtAz). Thus, all the com ponents o f the total lidar signal, Equations 2 .7-2.9, increase by a factor k (i.e. N s ^ kN S, N d ^ kN D, N b ^ kN B, N t o t ^ kN TOT) and the relative error decreases by a factor o f k 1/2
a n s ^ VkNs + kNB + kND _ Vn s + n b + n d _ a n s (2.13)
Ns kNs _ VkNS _ VkNS
Therefore, the raw photon count profiles can be acquired at high"resolution and then post" processed at low er resolution to increase the signal and decrease the relative error. For example, raw lidar profiles w ith a resolution o f At = 50 s and Az = 75 m can be post
processed at a resolution o f At = 30 min and Az = 1875 m, increasing the signal levels by a factor o f 900 and decreasing the relative error by a factor o f 30. This ability to m anipulate the resolution and statistical accuracy o f the lidar m easurem ents allows the operator to conduct the observations at high resolution and optim ize the perform ance o f the lidar system in real tim e and then post process the m easurem ents at low er resolution that yields a desired statistical accuracy. As the error depends on the product o f the tem poral and spatial resolution, so the resolution in tim e and space can change w hile m aintaining a constant product AtAz w ith no change in error. F or example, to have 1% errors at 62.5 km (Table 2.2) requires m easurem ents at a resolution o f 75 m and 39,200 s ( = 784,000 laser pulses/20 pps, 653 m in) or 2.94x106 m s. M easurem ents at a resolution o f 1 km and 2,940 s (49 m in) or a resolution o f 2 km and 1,470 s (24.5 m in) w ould also have 1% errors.
The quality o f the lidar signal can also be described by the signal-to-noise ratio (SNR), w hich is ju st the square o f the inverse o f the relative error,
( 2 1 4)
SNR
= ( 2 1 4 )A N s(z)2
On 2 -3 January 2012 the SNR is 3.0x105 at 42.5 km and decreases to 9 .8 x 103 at 62.5 km, 1.2 x 102 at 82.5 km, 1. 1x 101 at 92.5 km, and 7 . 1x 10-1 at 102.5 km.
W hile decreasing the m easurem ent resolution im proves the quality o f the lidar signal through increases in all com ponents o f the total signal, there are w ays to more dram atically increase the signal quality. For example, by increasing the atm ospheric signal, N s, w hile not increasing the background skylight signal, N b, and dark detector counts, N d . Increases in the laser pulse energy, E L, increase K S w hile leaving K B and K D unaltered. Increasing the repetition rate, R L, o f the laser increases all three signals ju st as it does in the case o f changes in m easurem ent resolution. Increases in the telescope area, A t , the receiver efficiency, r , and the atm ospheric transm ission, T, increase K S and K B w hile leaving K D unaltered. If the dark detector signal is negligible com pared to the