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Ice nucleating properties of soot, Kaolinite, & Goethite at conditions relevant for the lower troposphere Dymarska, Magdalena 2006

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ICE NUCLEATING PROPERTIES OF SOOT, KAOLINITE, & GOETHITE AT CONDITIONS RELEVANT FOR THE LOWER TROPOSPHERE  By Magdalena Dymarska B . Sc. Honours (Chemical Physics) University of Guelph, 2002  A THESIS S U B M I T T E D IN P A R T I A L F U L F I L L M E N T OF T H E R E Q U I R E M E N T S F O R T H E D E G R E E OF  M A S T E R OF S C I E N C E in T H E F A C U L T Y OF G R A D U A T E STUDIES CHEMISTRY  U N I V E R S I T Y OF BRITISH C O L U M B I A December 2005 © Magdalena Dymarska, 2005  A b s t r a c t  Although soot and mineral dust are both ubiquitous in the Earth's atmosphere, their contribution to the formation o f the ice phase in lower tropospheric clouds remains poorly understood. This thesis investigates the ice nucleating properties of a range of soot and mineral dust species in the deposition mode - below liquid water saturation - and at temperatures ranging from 239 to 258 K . It is found that under these experimental conditions soot exhibits no ice nucleating abilities whereas Kaolinite and Goethite are found to be effective ice nuclei. In experiments involving soot, ice nucleation was only observed at ~ 243 K on a few occasions. However, even at these temperatures the relative humidity with respect to ice (RHj) was close to water saturation when ice nucleation was observed, suggesting water nucleation may have occurred first followed by ice nucleation during the condensation process. The ice nucleating abilities o f soot exposed to atmospherically-relevant quantities o f ozone were also investigated. Even after an exposure equivalent to 80 ppb O3 at atmospheric pressure for 13.7 days, soot exhibited poor ice nucleating abilities. In the case o f Kaolinite and Goethite, ice particles consistently formed below ~ 252 K and at RHi well below water supersaturation, suggesting deposition nucleation is the dominant mode of ice formation under these experimental conditions. The ice nucleating abilities of soot and mineral dust were quantified by determining the heterogeneous nucleation rate coefficient, J ; heterogeneous nucleation theory was employed het  to determine the contact angle for formation of an ice germ on the surface of soot, Kaolinite, and Goethite. The experimentally-determined J  hel  was used to estimate the number density o f ice  particles that might form under given conditions of temperature, RHj, and particle number density for soot and mineral dust in the atmosphere. Whereas soot alone cannot account for the number density o f ice nuclei observed in field studies, these measurements show Kaolinite and Goethite may significantly contribution to the formation o f the ice phase in lower tropospheric clouds. This is consistent with recent field measurements and some recent laboratory data.  ii  Table of Contents  Abstract  ii  Table of contents  iii  List of Tables  vi  List of Figures  viii  Acknowledgements  xvi  Chapter 1  Introduction  1  1.1  Global climate forcing  1  1.2  The Earth' s atmosphere  2  1.3  Atmospheric nucleation processes  4  1.4  Tropospheric aerosols  6  1.4.1 Soot  6  1.4.2 Mineral dust  8  1.5  Thesis obj ectives  11  1.6  Thesis overview  11  Nucleation theory  13  2.1  Introduction  13  2.2  Saturation  14  2.3  Homogeneous nucleation of ice from supercooled water  16  2.4  Heterogeneous nucleation  19  2.4.1  19  Chapter 2  Chapter 3 3.1  Deposition nucleation - the stochastic model  Experimental technique  22  Introduction  22 iii  3.2  Sample preparation  22  3.2.1  Soot and clay specifics  22  3.2.1.1 Soot  22  3.2.1.2 Mineral dust  24  3.2.1.3 Or oxidized soot  24  3.2.2  Slide preparation  27  3.2.3  Sample preparation  28  3.3  Flow cell setup  3.4  Humidity control  3.5  29 •  30  3.4.1  RHj ramp experiments  31  3.4.2  Constant RHj experiments  32  Cell temperature & temperature calibration  32  Ice nucleating properties of soot  37  4.1  Introduction  37  4.2  Results  37  Chapter 4  4.2.1  RHj ramp experiments  37  4.2.2  Ice nucleation of soot particles oxidized by O3  39  4.2.3  JZ for RHj ramp experiments  41  4.2.4  Constant RHj experiments  41  4.3  Comparison with other results  44  4.4  Contact angle  47  4.5  Atmospheric implications  50  4.6  Conclusions  53  iv  Chapter 5  Ice nucleating properties of mineral dust: Kaolinite & Goethite  55  5.1  Introduction  55  5.2  Results  56  5.3  5.2.1  Ice nucleation on Kaolinite  56  5.2.2  Ice nucleation on Goethite  58  Comparison with other results  60  5.3.1  Kaolinite  60  5.3.2  Comparison of current data with previous mineral dust studies other than Kaolinite  62  5.4  Determination of onset J  5.5  Contact angle  70  5.6  Atmospheric implications  73  5.7  Conclusions  81  Summary and conclusions  82  6.1  Ice nucleating properties of soot and dust  82  6.2  The heterogeneous nucleation rate coefficient  84  6.3  Contact angle comparisons  84  6.4  Atmospheric implications  85  Chapter 6  References  het  for mineral dust  64  87  v  List o f Tables  Table 1.1:  A summary of commonly-observed mineral dust found in the troposphere along with their appropriate chemical formulas [Usher et al., 2003]  Table 3.1:  9  Physical characteristics of the soot types investigated in this study.  Data was  obtained from the manufacturer. 'Volatiles were determined by heating a sample in a muffle furnace for 7 minutes at 950°C.  2  B E T (Brunauer, Emmett, and  Teller)-Surface area was calculated from the N absorption isotherms recorded at 2  77 K  Table 4.1  23  Onset parameters pertaining to formation of a single ice crystal in multiple ramp RHi experiments for which J"£ was determined. The confidence level for J" el  p h et  is  95%. The limiting values of A G * , , r , and the contact angle are also included. C  AG* , was evaluated the assumption that the pre-exponential term, A, equals to C  10  26  cm^sec" . The size of the critical-size ice germ was evaluated based on Eq. 1  2.20 pertaining to the maximum Gibbs free energy of formation. Finally, a lower limit to the contact angle was determined using Eq. 4.4  Table 4.2  49  A summary of the number density of predicted ice particles per litre nucleated by soot particles for five locations: upper troposphere, free troposphere, U S A remote area, U S A urban-influenced rural area, and U S A polluted urban area. Elemental carbon mass concentrations vary with location and latitude.  Soot radius was  assumed to be 0.1 pm [Blake and Kato, 1995]. The experimentally-determined upper limit to the heterogeneous ice nucleation rate coefficient was 0.1 cm" sec 2  at 124% RHj and 247.5 K . The maximum amount of ice particles that can be produced for each scenario increases with soot density. Under these conditions Meyers et al. [1992] predicts that at 124% R H i the number of ice particles in the atmosphere is approximately 12 litre"  1  52  vi  Table 5.1  The heterogeneous nucleation rate coefficient for eleven separate Kaolinite experiments.  Parameters describing experimental conditions include onset  temperature and onset RHj at which either water droplets or ice crystals were observed to form first. The geometric surface area o f each Kaolinite sample is also included. The uncertainty in J  was evaluated by considering error in the  hel  observation time (20±10 sec) and the surface area o f particles. Here the lower and upper limits to J , are J'^, he  and J^,, respectively. The table also lists the  total number o f ice particles observed to form during the course of the experiment  Table 5.2  67  The heterogeneous nucleation rate coefficient experiments.  for six separate  Goethite  Parameters describing experimental conditions include onset  temperature and onset RHj at which either water droplets or ice crystals were observed to form first. The geometric surface area o f each Goethite sample is also included, noting that experiments no. 1, 5, and 6 were performed on the same sample.  The uncertainty in J  hel  was estimated by considering the error in the  observation time (20±10 sec) and the surface area o f the particles. Here the lower and upper limits to J  hel  are j'™ and J"£ , respectively. The table also includes the t  et  total number of ice particles formed during the course of the experiment  Table 5.3  69  Onset parameters pertaining to the formation of a single spherical-cap ice crystal in the Kaolinite and Goethite ramp RHj experiments for which Jh was et  determined. A G * , was evaluated by using the pre-exponential term, A, equal to C  10 cm" sec" . The size of the critical-size ice germ was evaluated by assuming no 26  2  1  lattice strain in Eq. 2.20. Finally, the contact angle was determined by using Eq. 4.4.  The uncertainty in J^,, AG *, r, and dxao was evaluated by considering act  uncertainty in the temperature, onset RHi, surface area, interfacial tension, and observation time (20±10 sec)  71  vii  List o f Figures  Figure 1.1  Estimated global, annual mean radiative forcing (Win' ) of a number of 2  anthropogenic and natural agents for the period from 1750 to 2000 [IPCC, 2001]. The height o f each rectangular bar denotes a central or best estimate value while its absence denotes no best estimate is possible. Each estimate is accompanied by an uncertainty range arising from the spread in published values of the forcing and physical understanding.  The "level of scientific understanding" index  illustrates the reliability of the forcing estimate and certainty in understanding of the impact of each variable on climate forcing  Figure 1.2  2  Temperature profde of the Earth's atmosphere as a function o f altitude, z (km) (from Wayne [2000]). The curve represents the mean structure for latitude 40°N during the month of June  Figure 1.3  3  Homogeneous and heterogeneous I N modes of activation. Supersaturation with respect to ice is defined as SS  5  ice  Figure 1.4  (a) A schematic model of microstructure of the diesel soot particle [Ishiguro, 1997]. This turbostratic, onion-like structure consists of short graphite segments, also known as crystallites, often grouped together in short stacks which are randomly rotated with respect to each other along the c-axis. (b) H R T E M image of ethanol-derived soot aggregates exhibits the alignment of graphite segments [Walet. al., 2003]  Figure 1.5  7  (a) A n electron micrograph showing large variability in mineral dust size and morphology [Falkovich et a.I, 2001]. (b) A n electron micrograph of Kaolinite, shows its highly crystalline structure (http://www.ktgeo.com/tEX.html). high  resolution  electron micrograph of Goethite  Schwertmann, 2003]  needles  (c) A  [Cornell  and 10  viii  Figure 2.1  Relative humidity with respect to ice (RHj  ce  = SV100%) for supercooled liquid  water in equilibrium with the surrounding water partial pressure using several different models (from Murphy and Koop [2005])  Figure 2.2  15  Equilibrium vapour pressure of water (dashed curve) [Koop et. al, 2000] and ice (solid curve) [Marti and Mauersberger, 1993] as a function of temperature  Figure 2.3  15  The Gibbs free energy, A G , for the formation of an ice cluster as a function of its size, r, and the saturation ratio, S. The energy of formation of an ice embryo is defined by Eq. 2.8. Under supersaturated conditions (i.e. S > 1) the ice germ will grow spontaneously only i f it reaches its critical size, r* (Eq. 2.9), and overcomes the energy barrier, AG* (Eq. 2.10). A n increase in the saturation ratio (S2 > Si) reduces the critical radius, r* , and the activation energy, AG* 2  Figure 3.1  1985]  27  Average size distribution of eight different samples o f Degussa F W 2 (Channel Type Black) deposited on hydrophobic glass slides  Figure 3.5  26  Image of a droplet positioned on top of a (a) clean glass slide and (b) a silanized glass slide  Figure 3.4  23  Sketch of the flow-tube coupled to the Chemical Ionization Mass Spectrometer (not to scale)  Figure 3.3  18  Two-dimensional model of n-hexane soot formed in a flame [Akhteretal,  Figure 3.2  2  28  (a) Three-dimensional cut away illustration of the flow cell, (b) Cross-section of the flow cell and location of microscope objective (Al=aluminium, P C T F E = polychlorotrifluoroethylene)  29  ix  Figure 3.6  Typical experimental trajectories of RHi, where temperature was reduced at a rate of 0.1 K m i n " , while the water partial pressure was constant. 1  The trajectories  were calculated using the saturation vapour pressure of water from Koop et al. [2000] and the saturation vapour pressure of ice from Marti and Mauersberger [1993]. The arrows show the experimental trajectory  Figure 3.7  31  This figure illustrates the change in surface area o f liquid droplets condensed on a hydrophobic slide in relation to RHj, dew point, R T D , and cell temperature (corrected for offset). As the cell temperature was ramped down (A) at a constant rate of 0.1 K min" , the RHj over deposited particles increased. Liquid droplets 1  were first condensed when R H  W  = 100% (B). They continued to grow in size (C)  until R H < 100% (D). Once again the cell temperature was decreased at a W  constant rate and the droplet surface area increased above 100% R H  Figure 3.8  W  35  Images were recorded as the temperature of the cell was ramped down and up at constant pmo- Labels (A) to (D) correspond to labels in Figure 3.7. The black deposits are Degussa FW2 soot on a hydrophobic surface. Each image is accompanied by experimental time, r n, and the RHi cc  Figure 4.1  36  The RHj at which liquid water droplets or ice particles were first observed as the RHj inside the cell was slowly increased. The open symbols indicate that water droplets were first observed (sometimes ice was observed to form after water droplets condensed) and the solid symbols indicate that only ice particles were observed with no indication of the formation of water droplets prior to ice formation. Panel (A) illustrates the results from a control experiment, where no particles were deposited on the hydrophobic substrate. Panel (B) to (F) illustrates the results from the six soot types listed in Table 3.1  Figure 4.2  38  The RHj at which water droplets were observed for Lamp Black 101 soot samples treated with a range of O3 exposures.  In all experiments water droplets were  observed first (i.e. ice particles did not form unless water droplets first  x  condensed).  The exposure times indicated in the plot are equivalent to an  atmospheric mixing ration of 80 ppb (see text for details)  Figure 4.3  40  (a) N-hexane (air/fuel = 2.4) soot particles on a hydrophobic slide. The surface area of the soot particles is 1.1-10 um . (b) Sample size distribution during the s  2  constant RHi experiment. The geometric mean diameter is 7.1 pm, with standard '  Figure 4.4  deviation of 1.2 pm  42  Change in relative humidity with respect to ice (black) during the constant RHj experiment. The liquid water saturation in terms of R H ; is also plotted (pink); it shows that water saturation was never reached. The drift in the frost point (green) and the cell temperature (blue) is also illustrated. For a period of approximately 8 hours cell temperature was held at 247.5 K and RHi = 124 ± 4%, close to water saturation. Water droplets or ice crystals were not observed to form.  Figure 4.5  43  A comparison of current results for Lamp Black 101 (open squares) with those of DeMott et al. [1999] (filled triangles).  Our data points correspond to the  conditions at which water droplets were observed using soot particles ranging in size from 1 to 40 um in diameter.  In these experiments water droplets were  always observed first. If ice did form it was only after the appearance of water droplets. The results from DeMott et al. correspond to the onset for which 1% of Lamp Black soot particles (a number mean diameter of 240 nm) nucleated ice. ..45  Figure 4.6  Lower limit to the contact angle vs. temperature. Here a critical-size, sphericalcap ice germ was assumed to form on an insoluble soot surface approximated to be flat. The contact angle was determined for N\ = 3.1-10 m" and o\ = 106.5 28  3  v  mJ m" using Eq.4.3 and Eq. 4.4. Plotted data corresponds to the constant and 2  ramp RHj experiments for which the soot surface area was known.  Since the  contact angle is a limiting value, it depends on liquid water saturation, and increases at lower temperatures  50  xi  Figure 5.1  The RHj at which liquid water droplets or ice particles were first observed to form on Kaolinite as the RHj inside the cell was slowly increased. The open symbols indicate that water droplets were first observed and the solid symbols indicate that only ice particles were observed with no indication of the formation of water droplets prior to ice formation  Figure 5.2  57  Images of ice crystals and Kaolinite particles from two different experiments at 245 and 243 K . Cell temperature was decreased at a constant rate of 0.1 K min" until ice crystals were observed to form  Figure 5.3  1  57  The RHj at which liquid water droplets or ice particles were first observed Goethite as the RHj inside the cell was slowly increased.  on The open symbols  indicate that water droplets were first observed and the solid symbols indicate that only ice particles were observed with no indication of the formation of water droplets prior to ice formation  Figure 5.4  59  Images of ice crystals and Goethite particles obtained from two different experiments at 241 and 239 K . Here cell temperature was decreased at a constant rate of 0.1 K min" until ice crystals were observed to form  59  1  Figure 5.5  A comparison of current onset RHj results for Kaolinite with those of Roberts and Hallett [1968] and BaUey and Hallett [2002]. Current data corresponds to the conditions at which water droplets (open squares) or ice crystals (solid squares) were observed. Roberts and Hallett [1968] observed immersion freezing above 254 K and deposition nucleation at lower temperature on particles between 0.5 to 3 pm in diameter.  The threshold of nucleation activity was taken as the  appearance of one ice crystal in ~10 particles. 4  The results from Bailey and  Hallett [2002] correspond to the onset of several ice crystals on Kaolinite particles (between 5 to 10 pm in diameter) adhered to a glass  filament  61  xii  Figure 5.6  A comparison of the current measurements of the onsets of ice nucleation of Kaolinite and Goethite particles with data for other mineral dust measured in previous work  Figure 5.7  53  (a) A n image of the bottom of the cell showing the deposition of ~ 300 particles of Kaolinite on a hydrophobic slide.  The surface area of the particles is  (6.5±0.8)-10 um . (b) Average size distribution of eleven different Kaolinite 4  2  samples. Geometric mean diameter is 8.0 um; geometric standard deviation is 1.2; surface area mean diameter is 19.7 pm (^parameters defined by Reist [1993])  Figure 5.8  66  (a) A n image of the bottom of the cell showing the deposition of ~ 1000 particles of Goethite on a hydrophobic slide. The surface area of the particles in view of the lens is (19±4)-10 pm . (b) Average size distribution of six different samples 4  2  of Goethite. Geometric mean diameter is 6.1 pm; geometric standard deviation is 1.2; surface mean diameter is 17.8 pm (parameters defined by Reist [1993])  Figure 5.9  68  Contact angle of an ice crystal formed under heterogeneous nucleation conditions vs. temperature.  Here a spherical-cap ice germ was assumed to form on a flat  insoluble Kaolinite or Goethite surface (geometric mean diameter equal to 8 pm and 6.1 pm) respectively. Neglecting lattice strain, the contact angle was determined for N\ = 3.1T0 m" and o\ = 106.5 mJ m" using Eq. 4.3 and Eq. 4.4. 28  3  2  v  Solid symbols correspond to experiments in which ice particles nucleated first, with no indication of the liquid phase prior to ice formation. Open symbols correspond to the limiting contact angles for experiments in which water droplets were observed to form first.  For experiments in which ice was observed to  nucleate first, the data is fitted to reflect the dependence of the contact angle on temperature. For Kaolinite and Goethite at temperatures between ~240 to 250 K, ^Kaolinite  =  -230.9 + 1.0 T and Ocoeihite = -276.7 + 1.2 T, respectively  72  xiii  Figure 5.10  The heterogeneous nucleation rate coefficient vs. supersaturation for Kaolinite. Each curve corresponds to a different temperature: 240, 245, and 250 K . contact  angle  at each  temperature  was determined  The  by applying the  parameterization Eq. 5.2. Assuming that the contact angle is independent of saturation, the nucleation rate coefficient (Eq 2.22) at each temperature is plotted as a function o f supersaturation. Each J  hel  is extrapolated up to R H  Dotted curves indicate the upper and lower bounds for J  hel  = 100%.  W  as a result of the  uncertainty in 9  74  Kaolinite  Figure 5.11  The heterogeneous nucleation rate coefficient vs. supersaturation for Goethite. Each curve corresponds to a different temperature: 240, 245, and 250 K . contact  angle  at each  temperature  was determined  The  by applying the  parameterization Eq. 5.3. Assuming that the contact angle is independent of saturation, the nucleation rate coefficient (Eq 2.22) at each temperature is plotted as a function o f supersaturation. Each J  hel  is extrapolated up to R H = 100%. W  Dotted curves indicate the upper and lower bounds for J  hel  as a result of the  uncertainty in 0  75  GoeMle  Figure 5.12  Number density o f ice vs. saturation for Kaolinite assuming mass density of dust in the atmosphere is 1 pg m" . The number density of ice is a function of the 3  nucleation rate coefficient determined i n Figure 5.10.  The solid curves  correspond to different temperatures of 240, 245, and 250 K for which the contact angle was 8.3°, 13.3°, and 18.3° respectively. Corresponding dotted curves reflect the standard deviation of 0  KaolMle  . The Meyers et al. [1992] parameterization  predicts the total number o f active I N in the atmosphere at a function of supersaturation, in the deposition mode and condensation freezing mode  Figure 5.13  78  Number density of ice vs. saturation for Goethite assuming mass density of dust in the atmosphere is 1 pg m" . The number density o f ice is a function of the 3  nucleation rate coefficient determined in Figure 5.11. The curves correspond to xiv  different temperatures of 240, 245, and 250 K for which the contact angle was 7.6°, 13.5°, and 19.4° respectively. standard deviation of # , , • T Goe  w  e  n e  Corresponding dotted curves reflect the  Meyers et al. [1992] parameterization predicts  the total number of active EST in the atmosphere at a function of supersaturation, in the deposition mode and condensation freezing mode  Figure 6.1  79  A comparison of the RHj at which water droplets or ice crystals were observed to form first for RHj ramp experiments involving soot, Kaolinite, and Goethite. Solid symbols correspond to experiments in which ice crystals were observed to nucleate first; open symbols correspond to experiments in which water droplets condensed  Figure 6.2  first  83  Contact angle as a function o f temperature for RHi ramp experiments with soot, Kaolinite, Goethite, and constant RHj experiment with soot.  Solid symbols  correspond to experiments in which ice crystals were observed to nucleate first; open symbols correspond to experiments in which water droplets condensed  xv  Acknowledgements  It has been my pleasure to cross paths with many incredible individuals who inspired and challenged me to think critically. This thesis is a tribute to the array of people who shaped my academic adventures.  First and foremost, I would like to thank Dr. Allan K . Bertram for  inspiring me to pursue atmospheric chemistry.  It has been a pleasure working under your  tutelage. Thank you for your patience and for all the conversations that clarified my thinking. A is for Amazing and Awesome! Y o u can now pat yourself on the back as your first one 'bites the dust'! Graduate school would not have been the same without the outstanding support of the entire Bertram group. In particular I would like to thank Dr. Ben J. Murray for keeping me on the right track over the past 2+ years; your advice on experimental design and subsequent analysis was invaluable. Thank you for being an outstanding mentor. And, yes, cubic ice rocks! A special thank you goes to Dr. Daniel A . Knopf for bombarding me with a million and one questions, and always keeping me on the edge of my scientific seat. I would like to express my gratitude to two Timbit junkies, Matthew T. Parsons and Jackson S. Mak, who continue to surprise me with their creative pranks to this day.  To Lori Anthony, Simone Gross, Sarah  Hanna, Jenna Riffell, Emily Simpson, and Dr. Pedro Campuzano-Jost for making the past two years a high point in my academic life. I will always remember our evenings at the Koerner's Pub, Storm the Wall, the Long Boat, tubing at Cypress, and endless B B Q s and parties which we have experienced together. The nuts and bolts, glass tubes, and research nuggets would not have been available without the handy work of the Mech shop and the Glass shop staff. I would particularly like to thank Des Lovrity, Raz Neagu, Oskar Greiner, Ken Love, and Brian Snapkauskas for solving my aperture dilemmas, and Brian Ditchburn for fixing my broken glass bits. Thanks for always keeping your doors open. I would also like to acknowledge Julie Michaud, Dr. Steve Cooke, Heidi Cooke, Christine Krumley, Dr. Mike Garry, and Angelo Ariganello for being my family away from home. To the sparkle in my life, Derek E . Brown, who is my strength, happiness, and a word of wisdom; I look forwards to the rest of our crazy lives together, be it scuba-diving deep below the ocean surf or sky-diving high above the clouds. I am most indebted to my parents who have sacrificed all that they knew in hope of finding a better life in Canada. You have provided me with the ability to take risks, think outside the box, and put a little bit of "heart and soul" in everything I do. Z calego serca, dziekuj? za wszystko. xvi  Chapter 1  Chapter I: Introduction, 1  Introduction  1.1  Global climate forcing  Scientific evidence strongly suggests that increased anthropogenic activity is influencing the Earth's atmosphere and climate [IPCC, 2001]. Measurements in the Arctic and Antarctic regions, for example, imply that the Earth's radiative budget is changing [Cavalieri, 1997]. Traditionally, radiative forcing is used to quantify the change in the Earth's radiative budget, where radiative forcing is defined as "a change in the net vertical irradiance at the tropopause due to an internal change, or a change in the external forcing of the climate system" [IPCC glossary, 2001]. Figure 1.1 illustrates the best estimate and uncertainty of the radiative forcing for a number of components of the climate system for the period 1750 to 2000. Some forcings, such as that by the greenhouse gases, are well understood; while other components of the climate system, such as the indirect aerosol effect, are very poorly understood. The indirect aerosol effect is driven by the ability of aerosols to act as cloud condensation nuclei ( C C N , Type 1) or ice nuclei (IN, Type 2). B y acting as C C N or IN, aerosols may impact microphysics, radiative properties, and the lifetime of clouds. In doing so, aerosols may change the ability of clouds to cool (Albedo Effect) and warm (Greenhouse effect) the atmosphere. Figure 1.1 only illustrates the estimated uncertainty in contribution of the indirect aerosol effect for Type 1 aerosols. The contribution of I N (Type 2) to radiative forcing is not included in the latest IPCC report because the confidence level in the quantitative estimates is very poor. However, the I P C C acknowledges that ice nuclei almost certainly play a critical role in the processes of mixed-phase and ice-phase clouds. Further research is clearly needed to understand the contribution of I N to climate change. To accurately predict the impact of the indirect aerosol effect, it is important to quantify the ice nucleating properties of atmospheric aerosols. For solid particles in the atmosphere, this can be achieved by determining the heterogeneous nucleation rate coefficient (defined as the number of nucleation events per unit of surface area and per unit of time). B y quantifying the ice nucleating properties of aerosols typically observed within the different regions of the atmosphere, the radiative forcing of IN may be better understood.  Chapter I: Introduction, 2  g>  Halocarbons N 0  2  Aerosols  2  CH,  E  2 e 3- I WN CD  CO  1  0  r Tropospheric ozone  •*  Black carbon from fossil fuel burning  Mineral Dust  a. *=  cn c  "a  o LL  > CO TJ DC  A  ^  Solar  Contrails cirrus  tn  |  Aviation-induced  0  Stratospheric ozone  o) •O = -1* o  Organic Sulphate J J J J  1 burning  fuel burning  Landuse (albedo) only  Aerosol indil ed effect  -2  High  Med.  Med.  Low  Very Low  Very Low  Very Very Low Low  Very Low  Very Very Very Low Low Low  Level of Scientific Understanding  Figure 1.1  Estimated global, annual mean radiative forcing (Wm~ ) of a number of 2  anthropogenic and natural agents for the period from 1750 to 2000 [IPCC, 2001]. The height of each rectangular bar denotes a central or best estimate value while its absence denotes no best estimate is possible. Each estimate is accompanied by an uncertainty range arising from the spread in published values of the forcing and physical understanding. The "level of scientific understanding"  index illustrates the reliability of the forcing estimate and certainty in  understanding of the impact of each variable on climate forcing.  1.2  The Earth's atmosphere  The Earth's atmosphere consists of the troposphere, stratosphere, mesosphere, and thermosphere (Figure 1.2). Each of these regions possesses its own but interrelated chemistry and physics. The boundaries between these layers are referred to as the tropopause, stratopause, and mesopause.  This thesis focuses on the ice nucleating abilities of aerosols at conditions  relevant to the lower troposphere.  Chapter 1: Introduction, 3 The troposphere extends up to 8 to 18 km in altitude with its altitude dependent on latitude and season [Wayne, 2000]. The troposphere is the densest part of the atmosphere (up to 80% of the total mass of the atmosphere), with temperature decreasing at higher altitudes up to the tropopause. The temperature of the tropopause can be as low as 180 K .  T(K) Figure 1.2  Temperature profile of the Earth's atmosphere as a function of altitude, z (km)  (from Wayne [2000]). The curve represents the mean structure for latitude 40°N during the month of June.  Chapter 1: Introduction, 4 1.3  Atmospheric nucleation processes  Clouds cover approximately 60% of the Earth's atmosphere with only 10% of the total generating precipitation [Seinfeld and Pandis, 1998].  In the troposphere, they are created  through two mechanisms of isobaric and adiabatic cooling [Seinfeld and Pandis, 1998]. Isobaric cooling refers to the cooling of an air parcel under constant pressure. This may involve radiative losses of energy or a horizontal movement of an air mass over a colder surface. Under adiabatic conditions an air parcel rises vertically and cools without releasing heat to the surroundings during its expansion. In the atmosphere, upward air flow may be caused by solar heating of the Earth's surface, a cold front forcing a warm air mass aloft, or a mountain range at an angle to the wind. The eventual formation of the liquid or the ice phase is a result of homogeneous or heterogeneous nucleation. Homogeneous ice nucleation occurs in a supersaturated water vapour phase and is defined as the formation of ice from the gas or the liquid phase, in the absence of foreign substances [Seinfeld and Pandis, 1998]. Homogeneous ice nucleation from the vapour phase is unlikely to occur under normal atmospheric conditions because supersaturations  are never large enough [Young, 1993].  atmospheric  Heterogeneous ice nucleation  corresponds to the nucleation of an ice germ on a foreign substance, such as an ion or a solid particle [Seinfeld and Pandis, 1998]. Four modes of heterogeneous ice nucleation exist: deposition nucleation, condensation freezing, contact freezing, and immersion freezing [Pruppacher andKlett 1997; Vali 1985]. The various modes of ice phase formation are illustrated in Figure 1.3. Deposition nucleation occurs when vapour adsorbs onto a solid surface and is transformed into ice below water saturation. Condensation freezing refers to the sequence of events whereby cloud condensation initiates freezing of the condensate.  Immersion freezing occurs when ice nucleates on a solid particle  immersed in a liquid droplet, and contact freezing occurs when a solid particle collides with a liquid droplet, resulting in ice nucleation [Vali, 1985; Pruppacher and Klett, 1997].  Often  condensation freezing and immersion freezing are grouped together due to the similarity in these modes. The theoretical treatments for the mechanisms relevant to this thesis will be discussed in Chapter 2. At temperatures between 273 and ~235 K ice formation necessarily occurs though one of these heterogeneous mechanisms.  This temperature range is most relevant to the lower  Chapter 1: Introduction, 5 troposphere [Seinfeld and Pandis, 1998].  However, much disagreement on the dominant  heterogeneous mode of formation of ice in the atmosphere exists particularly because very little is known about the ice nucleating properties of insoluble aerosols such as soot and mineral dust.  Homogeneous mechanism  Heterogeneous mechanisms Deposition nucleation Condensation freezing Immersion freezing Contact freezing  Temperature [, time f Legend:  "  '  •  C C N , soluble particle IN, insoluble particle Liquid phase Ice phase  Figure 1.3:  Homogeneous and heterogeneous I N modes of activation. Supersaturation with  respect to ice is defined as SS . ice  Chapter I: Introduction, 6 1.4  Tropospheric aerosols  Aerosols in the Earth's atmosphere are relatively stable suspensions of liquid or solid particles in the gas phase and range in size from 0.002 to 200 pm in diameter [Finleyson-Pitts and Pitts Jr., 2000]. Four distinct aerosol size classifications are recognized. Particles larger than 2.5 pm in diameter are termed course particles, whereas those smaller than 2.5 pm in diameter are fine particulates. The majority of aerosols found in the troposphere are grouped under the fine mode which can be further subdivided into three categories: accumulation mode (-0.08 - 2.5 pm), transient or Aitken mode (0.01 - 0.08 pm), and ultrafme mode (<0.01 pm). Aerosols enter the troposphere by means of natural, anthropogenic, direct, and secondary sources. Direct sources include biomass burning (54 Tg/yr), sea spray loading (54 Tg/yr), and fossil fuel combustion (29 Tg/yr) [IPCC, 2001]. season, location, and altitude.  Aerosol composition varies strongly with  A n average composition of urban fine particles, based on  measurements at several different sites, is 31% organic carbon, 28% sulfate, 9% elemental carbon, 8% ammonium, 6% nitrate, and 18% other species [Heintzenberg, 1989]. Whereas the homogeneous ice nucleating properties of solution droplets composed o f H , +  are fairly well understood [Martin, heterogeneous  NH*, SO ', NO~ 1  2000; Koop et al., 2000], further research on the  nucleating behaviour of insoluble soot and mineral dust aerosols under  tropospheric conditions is still needed.  1.4.1  Soot  Soot is ubiquitous in the atmosphere.  In fact, ice core measurements show that soot  concentrations have increased from pre-industrialized to modern times [Lavanchy et al, 1999]. Globally, 13 Tg o f soot is emitted into the troposphere annually, 54-57% of which is a result of fossil fuel combustion [Cooke and Wilson, 1996]. With an onion-like structure illustrated in Figure 1.4 (a,b), soot results from incomplete combustion processes and consists of an elemental carbon (EC) inner core coated with a layer of a polycyclic aromatic hydrocarbons (PAHs) and aliphatics under an amorphous outer shell of volatile compounds [Wal and Tomasek, 2003; Steiner et al, 1992; Haynes and Wagner, 1981; Finlayson-Pitts and Pitts Jr., 2001].  Chapter 1: Introduction, 7  Figure 1.4:  (a) A schematic model of the microstructure of the diesel soot particle [Ishiguro,  1997]. This turbostratic, onion-like structure consists of short graphite segments, also known as crystallites, often grouped together in short stacks which are randomly rotated with respect to each other along the c-axis. (b) H R T E M image of ethanol-derived soot aggregates exhibits the alignment of graphite segments [Wal and Tomasek, 2003].  If soot are effective IN, they have the potential to significantly impact the Earth's climate indirectly by changing the properties and lifecycle of mixed-phase and ice clouds on a global scale [DeMott, 2002; DeMott et al, 1997; Gierens, 2003; Jensen and Toon, 1997; Lohmann, 2002; Lohmann and Feichter, 2005]. In the lower troposphere, an increase in soot particles may lead to more frequent glaciation of supercooled clouds. The presence of insoluble soot particles may increase the amount of precipitation by means of the ice phase. This may further reduce the cloud cover in the lower troposphere and result in increased absorption of solar radiation [Lohmann, 2002; Lohmann and Feichter, 2005]. However, this indirect aerosol effect on climate is highly uncertain, in part, because the conditions at which ice nucleates on soot particles in the atmosphere are poorly quantified [Penner et o/.,2001]. At present there have only been three studies on the ice nucleating ability of soot particles at temperatures above 238 K [DeMott, 1990; Diehl and Mitra, 1998; Gorbunov et al, 2001]. The soot particles investigated in the previous studies were produced from the combustion of acetylene [DeMott, 1990], kerosene [Diehl and Mitra, 1998], benzene and toluene [Gorbunov et al, 2001], as well as soot produced by thermal decomposition of benzene [Gorbunov et al, 2001].  These measurements suggest that soot particles are potentially important IN in the  atmosphere.  However, more work is still needed to understand ice nucleation on soot at  temperatures relevant for the lower troposphere.  For example, the ice nucleating ability as a  Chapter 1: Introduction, 8 function of relative humidity (RH) needs to be investigated, since the previous studies mainly focused on ice nucleation at or slightly above liquid water saturation. Further complications arise from the fact that soot is exposed to tropospheric oxidants such as hydroxyl and nitrate radicals, and ozone [Mohler et al., 2005a; Lelievre et al., 2004] which may change soot's nucleating abilities. Although previous authors [Gorbunov et al., 2001; Mohler et al., 2005a; Lelievre et al., 2004] speculate that soot's abilities to nucleate ice should improve following exposure to atmospheric oxidant, quantitative research is yet to be presented. In this thesis the ice nucleating properties of soot in the deposition mode and at conditions relevant for the lower troposphere are considered.  1.4.2  M i n e r a l dust  Mineral dust is one of the most abundant aerosol species in the troposphere with global source strength estimates ranging from 1000 to 5000 Tg/yr, o f which up to 50% may be anthropogenic in origin [IPCC, 2001; Raes, 2000].  90% of global airborne mineral dust is  generated in the northern hemisphere where it is deposited [Usher et al., 2003]. Typical volume median diameters of particles are of the order of 2 to 4 pm [IPCC 2001]. The two primary sources of dust are the deserts of the Mongolia and Saharan regions [Perry et al, 1997; Prospero 1996]. Additional sources include dry lake beds and regions dried as a result of anthropogenic activity [IPCC, 2001]. Multiple studies of the elemental component of windblown dust (originating from various locations around the world) report that mineral dust is approximately 60% SiC>2 and 10-15% AI2O3  [Usher et al, 2003]. The percentage of other oxides such as Fe203, M g O , and CaO, are  slightly more varied and dependent on source location [Usher et al, 2003]]. Some of the most common minerals found in atmospheric dust are listed in Table 1.1. Typically Saharan dust is dominated with mica-illite (55 to 65%), quartz (14 to 20%), kaolinite, hematite, chlorite, and calcite (all of which have concentrations less than 15% of the total) [Glaccum and Prospero, 1980;  Usher et al, 2003]. The presence of Kaolinite in atmospheric dust has been noted in  several studies [Goudie and Middleton, 2001; Shi et al, 2005]. Clay samples in Tamanrassett, Sessali, and In Guezzam, Africa, have been found to contain up to 25% Kaolinite dust [Goudie and Middleton, 2001]. Iron oxides have been observed in the atmosphere [Weber et al, 2000;  Chapter 1: Introduction, 9 Table 1.1:  A summary of commonly-observed mineral dust found in the troposphere along  with their chemical formulas [Usher et al, 2003].  Mineral  Formula  Calcite Chlorite Corundum Dolomite  CaC0 A .6Z Oio(OH) a-Al 0 CaMg(C0 )  Feldspars  WZ4O3*  Gypsum Halite Hematite Illite Kaolinite Magnesite Montmorillonite Mica Opal Palygorskite Quartz  CaS0 -2H 0 NaCl oc-Fe 0 (K,H 0)(Al M Fe) (Si,Al)40 o[(OH) H 0] Al Si Oio(OH)8 MgC0 (Na,Ca) . (Al,Mg) Si Oio(OH)2-nH20 W(X,Y) . Z Oi (OH,F) Si0 nH 0 (Mg,Ai) Si4Oi (OH)-4H O Si0  3  5  4  2  3  3  4  )  g)  3  2  4  2  2  2  3  a 8  1  2)  2  4  3  0  33  2  4  e  2  3  4  0  2  2  2  2  0  2  2  "Typically A = Al, Fe, Li, Mg, Mn, and/or Ba; Z = A1,B, Si, and/or Fe. ^Typically W= Na, K, Ca, and/or Ba; Z = Si and/or Al. "Typically W = K or Na; X and Y = Al, Mg, Fe2+, Fe3+, and Li; Z = Si and A l .  Hoffinann et al, 1996; Penn et al, 2001; Behra and Sigg, 1990]. Hematite, a - F e 0 , occurs in 2  3  association with Goethite, a-FeOOH [Cornell and Schwertmann, 2003]. B y mass, Hematite is typically present below 5% in mineral dust; it is often morphologically present as coatings on clay grains [Hung et al, 2003; Penn et al, 2001]. Goethite is the most common iron oxide in soils [Cornell and Schwertmann, 2003]. As is shown in Figure 1.5 (a), a typical atmospheric sample of mineral dust may contain a variety of sizes and morphological characteristics. In addition, during long-range transport the particles may undergo physical and chemical changes which may alter their ice nucleating abilities [Posfai et al 1994; Marin, 2000; Prospero, 1999; Usher et al, 2003]. Figure 1.5 (b) shows an S E M image of a pure Kaolinite sample and Figure 1.5 (c) illustrates Goethite microstructure. The ice nucleating abilities of mineral dust in the atmosphere remain poorly understood. Previous research suggests that mineral dusts found in the Earth's atmosphere exhibit strong ice nucleating abilities [Roberts and Hallett, 1968; Bailey and Hallett, 2002; Zuberi et al, 2002;  Chapter 1: Introduction, 10 Hung et al, 2003; Archuleta et al., 2005; Mohler et al., 2005b; Knopf and Koop, submitted]. Consequently, these results suggest that dust is potentially a very important I N and may contribute strongly to the indirect aerosol effect  [Lohmann and Feichter, 2005]. Several studies  have attempted to quantify the nucleating properties of mineral dust in immersion mode  [DeMott, 2002; Hung et al., 2003; Archuleta et al., 2005] and in the deposition mode [Roberts and Hallett, 1968; Bailey and Hallett, 2002; Archuleta et al., 2005]. However, further research is necessary to accurately predict the contribution of mineral dust to the formation of ice at lower tropospheric conditions.  This thesis considers the ice nucleating properties of Kaolinite and  Goethite in the deposition mode at conditions relevant to the lower troposphere. (a)  Figure 1.5: morphology  (a) A n electron micrograph showing large variability in mineral dust size and  [Falkovich et a.I, 2001].  (b)  An  electron  micrograph  shows its highly crystalline structure (http://www.ktgeo.com/tEX.html). electron micrograph of Goethite needles  of  Kaolinite,  (c) A high resolution  [Cornell and Schwertmann, 2003].  1.5  Chapter 1: Introduction, 11  Thesis objectives  The primary goal of this thesis is to investigate the ice nucleating properties of soot, Kaolinite, and Goethite at conditions relevant for the lower troposphere. The thesis focuses on ice nucleating abilities of soot and mineral dust particles in the deposition nucleating mode below water saturation - and at temperatures ranging from 239 to 258 K. Experiments were done as a function of both temperature and relative humidity. For these studies several different types of soot and black carbon particles with a range of physical and chemical properties were used. The ice nucleating abilities of Kaolinite and Goethite were investigated using similar experimental parameters as the experiments involving soot. Additionally, experiments were carried out to determine if the oxidization of soot particles increases their ability to act as ice nuclei. It has previously been speculated that oxidation of soot by ozone in the atmosphere will increase the ice nucleating ability of soot particles [Gorbunov et al, 2001]. As a test of this hypothesis, Lamp Black 101, a commercial black carbon, was exposed to ozone for extended periods of time, and then tested for its ice nucleating ability in the deposition mode. 1.6  Thesis overview  This thesis is organized into six chapters. Chapter 1 covered background information. This included an overview of the indirect aerosol effect, tropospheric nucleation modes, and the current state of knowledge of the contribution of soot and mineral dust to ice formation. The theoretical treatment of homogeneous and heterogeneous ice nucleation is discussed in Chapter 2. The apparatus and the methods used in this thesis are described in Chapter 3. Special attention is paid to sample preparation, flow cell design, and humidity control. Two separate experimental methods are discussed: ramp RHj experiments, constant RHi experiments. Experimental requirements for oxidation of soot with atmospherically-relevant quantities of ozone are also presented.  The results for several different types of soot and black carbon  particles with a range of physical and chemical properties (including soot exposed to ozone) are presented in Chapter 4. The ice nucleating ability of soot is quantified by determining the upper limit to the heterogeneous nucleation rate coefficient, J , which is used to calculate the lower u p  h el  limit to the contact angle for the formation of a single ice germ on the surface of soot. Atmospheric implications of these results are considered by using J"  P h t  values to determine the  Chapter 1: Introduction, 12 maximum number density of ice crystals that can be nucleated by soot at conditions relevant to the lower troposphere. By comparing this value against parameterizations of ice nuclei in the atmosphere [Meyers et al., 1992], soot is shown not to be an important IN in the lower troposphere. Results of the RH ramp experiments for Kaolinite and Goethite are presented in t  Chapter 5. The ice nucleating properties of these mineral dust species are also quantified by determining the heterogeneous nucleation rate coefficient, Jh , which is used to calculate the et  contact angle of an ice embryo formed from supercooled water vapour. By utilizing the contact angle, the atmospheric implications of these mineral dust results are considered. Both Kaolinite and Goethite are found to exhibit strong ice nucleating properties which implies this type of mineral dust may significantly alter the formation of the ice-phase in lower tropospheric clouds. The thesis closes with Chapter 6 in which comparisons between the ice nucleating properties of soot and mineral dust are drawn.  Chapter 2: Nucleation theory, 13  Chapter 2  Nucleation theory  2.1  Introduction  Ice can form from supercooled water vapour and supercooled water droplets through various modes of homogeneous or heterogeneous nucleation. In order to place these theoretical concepts in an atmospherically-relevant context, variables relevant to formation of water and ice in the atmosphere are defined first. This includes the definition of the partial pressure of water and ice as well as saturation ratios.  The theoretical approaches to homogeneous and  heterogeneous ice-forming mechanisms are reviewed in sequence.  While homogeneous  nucleation is not directly related to the experimental study, the theory introduces concepts relevant to heterogeneous ice nucleation. Homogeneous nucleation of a single ice crystal from a supercooled liquid droplet is discussed in terms of the Classical Nucleation Theory (CNT). The equations presented here are taken from Young [1993]. CNT is used to determine the activation energy for the formation of a critical-size cluster from the homogeneous nucleation rate coefficient, J},  .  om  These principles are  applied to the formation of an ice germ on a foreign and insoluble particle. Through a simple approach, the heterogeneous nucleation rate coefficient, Jf, , for the formation of a spherical-cap et  ice embryo on a flat and insoluble surface is determined. temperature, saturation, and the contact angle is assessed.  The relationship between  Jf, , et  Chapter 2: Nucleation theory, 14 2.2  Saturation  Formation of a liquid droplet or an ice crystal in the atmosphere depends on saturation and temperature.  The saturation ratio with respect to liquid water (S ) and ice (S.) can be w  determined from the equilibrium vapour pressure:  S... = PHIO  Eq. 2.1  P (T) W  S, = PHIO  Eq. 2.2  P,(T)  where pmo is the water vapour partial pressure, p (T) is the saturation vapour pressure over pure w  liquid water, and pi(T) is the saturation vapour pressure over ice. formulas have been developed to describe  A number of empirical  p (T) and p (T) [Murphy and Koop, 2005]. The t  w  discrepancy between some of these models is illustrated in Figure 2.1 for p,. For the purpose of this study, inconsistencies in defining the partial pressure of water and ice between 253 and 239 K are very small and consequently will not alter the conclusions drawn from results presented in this thesis. For the purpose of this study the partial pressure of ice and water is defined by the following equations:  p (T) = exp 28.868--  6132.9^  Eq. 2.3  i  and (  210368 +131.4387/ -  P (T) = P-,(T)^V W  3323730 J  U  ^ - 41729.1 ln(7/)  RT  Eq. 2.4  where the units of pressure and temperature are Pa and Kelvin respectively. R is the molar gas constant (8.31447 J mol" K" ). 1  1  The equation for the partial pressure of ice was obtained by  Marti and Mauersberger [1993] and is based on direct measurements. The equation for the partial pressure of liquid water was proposed by Koop  et. al. [2000]. Consequently, since p(T)\ is  lower than p (T) for a given temperature (Figure 2.2), Si will always be higher than S for a W  given PH2O(T) below the freezing point.  w  Chapter 2: Nucleation theory, 15  180 Figure 2.1  200 220 Temperature [K]  Relative humidity with respect to ice ( R H i  ce  240  = Sr 100%) for supercooled liquid  water in equilibrium with the surrounding water partial pressure using several different models (from  Murphy and Koop [2005]). H  1 e  j  1—i  1  ,  ,  ,  ,  .,  -1—i—I—r—|—i—|^—r—i—i—,—i 200 210 220 230; 240 250  r  —,  ,  r  | — i — p i 260 1270  Temperature / K  Figure 2.2: (solid curve)  Equilibrium vapour pressure of water (dashed curve) [Koop et. al, 2000] and ice  [Marti and Mauersberger, 1993] as a function of temperature.  Chapter 2: Nucleation theory, 16 2.3  Homogeneous nucleation of ice from supercooled water  In Classical Nucleation Theory the change in Gibbs free energy associated with the formation of a single spherical solid cluster of / molecules formed in pure water is defined as AG,- = Kpsolid ~P ) liq  + Acr  Eq. 2.5  LI  where p nd and pu are the chemical potentials of the solid and liquid phase, and au corresponds so  q  to the interfacial tension between the liquid water and the ice phase over the surface area, A. The number of molecules, /, within the cluster is defined as  • 4*r?N,  i =  ^  — 3  _  Eq. 2.6 M  where n is the radius and TV) is the molecular concentration of a spherical crystalline cluster. The chemical potential term in Eq. 2.5 can be expressed in terms of the saturation ratio: - (Msoiid ~ MIUM) = kT\n S,  Eq. 2.7  where k is the Boltzmann constant and T is absolute temperature. Using the above definitions and assuming that A = 4n r , Eq. 2.5 can be rewritten as 2  t  47zr jV 3  AG, =  i - W l n S , + 4m-, a 2  LI  Eq. 2.8  The change in Gibbs free energy of formation consists of two terms. The first term in Eq. 2.8 describes the decrease in Gibbs free energy due to the transfer of molecules from the supercooled liquid to the solid phase (i.e. p nd < Puqutd)- The second term corresponds to the amount of work so  required to create an interface between the two phases. The change in Gibbs free energy of formation of a solid cluster as a function of its radius and saturation is shown in Figure 2.3. When the conditions are under-saturated (S < 1) embryos will never reach a critical size because they are energetically not favoured (both terms in Eq. 2.8 are positive). Since the particles do not reach their critical size they are energetically unstable and the liquid will not freeze.  Under supersaturated conditions (Si > 1) embryos with radii  smaller than the critical size (r') will also dissociate because they are energetically not favoured. However, i f the cluster reaches its critical size, the first term in Eq. 2.8 will be larger than the surface energy term, allowing the cluster to grow continuously. This critical radius is determined by setting the derivative of the activation energy to zero and solving for r* to give: •  2cr,,  Chapter 2: Nucleation theory, 17 Consequently, the barrier for formation of a stable cluster of radius r is defined by the following parameters: , A  =—,  G  16;rcr?,  —  E  q  .  2  .  32na],  i =3(iV,A:rin5,.)  1  0  Eq.2.11  t  2  where AG* is the barrier to formation of a critical germ and i is the critical number of molecules needed to achieve a single ice nucleation event. A n increase in the saturation ratio (S2 > Si) reduces r , AG*, and 1 . The experimental quantity of interest is the rate at which ice germs appear in the system as a function of the saturation ratio and temperature. This frequency of nucleation, termed the homogeneous nucleation rate coefficient  (Jhom),  is measured as the number of ice crystals  appearing per unit volume and per unit time: J*,™ - A  hom  where Ah  om  f AG > exp kT is the pre-exponential factor approximated to 1 T 0  AG , = AG* + AG . ac  djff  Eq. 2.12 28  cm" sec"' [Young, 1993] and 3  In homogeneous freezing the activation energy for the formation of a  critical size embryo in the liquid phase depends on AG* and the molar Gibbs free energy of activation for the diffusion of molecules across the liquid-solid boundary, A G ^ .  Chapter 2: Nucleation theory, 18  Figure  23:  The Gibbs free energy, AG, for the formation of an ice cluster as a function of its  size, r, and the saturation ratio, S. The energy of formation of an ice embryo is defined by Eq. 2.8. When S < 1, ice nucleation is not favoured. Under supersaturated conditions (i.e. S > 1) the ice germ will grow spontaneously only if it reaches its critical size, r* (Eq. 2.9), and overcomes the energy barrier, AG* (Eq. 2.10). An increase in the saturation ratio (S > Si) reduces the 2  critical radius, r *, and the activation energy, AG*. 2  Chapter 2: Nucleation theory, 19 2.4  Heterogeneous nucleation When the relative humidity is not high enough or the temperature is not low enough for  homogeneous nucleation to occur, ice formation is attributed to the presence of solid insoluble particles which may promote ice formation. However, several requirements must be met before such a solid acts as an I N and can induce ice nucleation [Pruppacher and Klett, 1997]. First, a particulate acting as an IN must be partially-soluble or insoluble. If it is soluble, it will dissolve into the aqueous phase and will not provide a rigid substrate surface needed for ice germ formation.  Second, since the ice crystal lattice is held by hydrogen bonds, I N  nucleating properties will be enhanced i f they contain hydrogen bonds at the particle surface. However, any difference in the geometric arrangement of these bonds between the substrate and the ice phase may lead to elastic strain, e, within the ice lattice. Elastic strain increases the activation energy of formation of an ice germ. If the misfit between the substrate lattice and the ice lattice is small, the ice lattice will have coherent geometry and nucleate on the substrate surface; however, i f the misfit is large, ice will not form on the substrate surface. Heterogeneous nucleation may be considered either in terms of the singular hypothesis.  stochastic or the  The stochastic model follows the homogeneous classical point of view  where the effects of I N depend on the probability of the random nucleating events. The singular model recognizes individual nucleating characteristics of active sites located on the surface of the IN. Presence of the so-called active sites may allow formation of crystalline germs at characteristic temperatures. Thus, the density of active sites available for ice formation at a specific temperature determines nucleation efficiency. In this work, the stochastic model will be employed to quantify the ice nucleating abilities of soot and mineral dust.  2.4.1  Deposition nucleation - the stochastic model The stochastic model of deposition nucleation assumes that at a given temperature all  equal-sized embryos have an equal probability of reaching their critical size. Consequently, the basic theory for deposition nucleation on a uniform and insoluble surface is an extension of the homogeneous nucleation approach. The equations presented here are taken from Young [1993] and describe heterogeneous nucleation of ice from the vapour phase on a flat insoluble surface.  Chapter 2: Nucleation theory, 20 More complex treatments involve spherical aerosols which will not be discussed here (see Young [1993] for further discussion). The expression for the thermodynamic potential of an ice embryo (with the shape of a spherical cap) formed on a flat substrate resembles Eq. 2.8, but accounts for the elastic strain, e, in the bulk ice lattice as well as the ice-substrate interface energy,  AG, = V (-N kT\nS +C e )+<r A  +{cr  2  i  l  i  lv  iy  Eq. 2.13  IN  In the above expression C is the estimated strain coefficient of 1.7x10' dynes cm" at 273.15K 1  2  [Turnbull and Vonnegut, 1952], e is the elastic strain, TV, is the molecular concentration of the ice embryo, and  CJVN, 0W,  and  CT/K  correspond to the vapour-nucleus, ice-nucleus, and ice-vapour  interfaces. Aw and AJN correspond to the areas of the ice-vapour and ice-nucleus interfaces, and Vi is the total volume of the spherical cap ice embryo. These variables are defined geometrically as 2  A y=2nr (l-m)  Eq. 2.14  w =w (l-m )  Eq. 2.15  I  l  A  2  2  77T* ^  K =-^-(2 + w)(l-w)  Eq.2.16  2  <  where m = cosf? = — — —  Eq. 2.17  Once these geometric terms are substituted into Eq. 2.13, the free energy of formation of a single ice germ over the flat insoluble surface simplifies to AG,=/(m)  —'- (- N kT ln S, + C e )+ 4nr cr 2  t  2  n  Eq. 2.18  where ~\2  (2 + w X l - w ) f(m) = ± ^ '-  2  Eq.2.19  The geometric term introduces the contact angle, 6, which controls the ice nucleating ability of the aerosol. If 0< 180°,/(w) < 1 and heterogeneous nucleation may take place. However, i f 9= 180°, homogeneous condensation w i l l commence in the gas phase once a germ of a critical radius, r, forms at a sufficiently high supersaturation.  Chapter 2:'Nucleationtheory, 21 Assuming heterogeneous nucleation takes place, the critical radius for spontaneous growth of the ice germ formed from vapour phase is consequently defined as 2(7,y  r =N,kT]nS* •  Eq. 2.20  -Ce  2  The heterogeneous nucleation rate coefficient for the formation of a critical-size ice embryo is -AG* aa  he, = Aexp J  -\6na] f(m)  = A exp  v  '[ikT^fiTlnS,-Ce )  Eq. 2.21  2 2  kT  A is often estimated to be equal to {lO^Jcm^sec" [Pruppacher and Klett, 1997]. The particulate 1  radius is approximated to a flat surface on which the ice germ forms. Assuming the ice embryo is independent of lattice strain, the above equation is often written in the following form  J =Aexp hel  r  3  (lnS,.)  [Vali, 1999] Eq. 2.22  2  where  B=  \67tf(m) 3N?k  Eq. 2.23  3  The equation above shows that J/, , can change rapidly with saturation, temperature, and surface e  tension. B y applying these theoretical equations to experimental observations the contact angle of soot, Kaolinite, and Goethite will be quantified in this thesis.  Chapter 3: Experimental technique, 22  Chapter 3  Experimental technique  3.1  Introduction The apparatus used in these studies consisted of an optical microscope coupled to a flow  cell in which the relative humidity could be accurately controlled. In the current experiments soot or mineral dust particles were deposited on the bottom surface of the flow cell; the relative humidity with respect to ice (RHj) inside the cell was increased, and the conditions under which water droplets or ice crystals formed were determined with a reflected-light microscope.  3.2  Sample preparation  3.2.1  Soot and clay specifics  3.2.1.1  Soot The various soots and black carbon used in these studies are listed in Table 3.1. Three  different samples of n-hexane soot were provided by Dwight M . Smith (University of Denver). The first sample was produced by burning n-hexane under ambient conditions in an open vessel, resulting in a diffusion flame. The second and third samples were generated using an apparatus designed for producing soot from flames with variable oxygen to fuel ratios. It has been shown that there is a linear relationship between the state of soot surface oxidation and the air to fuel ratio  [Chughtai et al., 2002]. The International Steering Committee for Black Carbon Reference  Materials has recommended using n-hexane soot as a model for soot in the atmosphere because a large amount of soot characteristics and reactivity data already exists in the scientific literature on this type of soot and because of the option to vary the n-hexane soot properties by varying the combustion conditions (http://www.du.edu/~dwismith/bcsteer.html). Properties of n-hexane soot have been documented by Smith and co-workers [Chughtai et al., 2002; Akhter et al., 1985]. A two-dimensional model of n-hexane soot as formed in a flame is shown in Figure 3.1.  Chapter 3: Experimental technique, 23 Volatiles  1  Type of Soot  (%)  BET Surface Area  2  c:%  H %  o%  N%  N-hexane soot - diffusion flame  n/a  89±2  87 to 95  1.6 to 1.2  11 to 6  n/a  N-hexane soot - Air/Fuel=0.53  n/a  100+2  n/a  n/a  n/a  n/a  N-hexane soot - Air/Fuel=2.4  n/a  156+11  n/a  n/a  n/a  n/a  Lamp Black 101  1  20  98.5  0.4  0.4  0.1  Degussa FW2 (Channel Black)  17  460  88  1.1  9.9  0.7  Printex 40 (Furnace Black)  0.9  90  n/a  n/a  n/a  n/a  Table 3.1:  Physical characteristics of the soot types investigated in this study.  obtained from the manufacturer. furnace for 7 minutes at 950°C.  Data was  Volatiles were determined by heating a sample in a muffle 2  B E T (Brunauer, Emmett, and Teller)-Surface area was  calculated from the N absorption isotherms recorded at 77 K . 2  Figure 3.1:  Two-dimensional model of n-hexane soot formed in a flame  [Akhter et al., 1985].  Chapter 3: Experimental technique, 24 Lamp Black 101, Degussa FW2 (which is a channel type black), and Printex 40 (which is a furnace type black) are commercially available black carbon. Degussa F W 2 is post-treated with NO2 and has been used in the past in laboratory heterogeneous chemistry studies (see for. example  [Choi and Leu, 1998; Disselkamp et al, 2000; Tabor et al, 1994]). Lamp Black 101 is  essentially non-volatile at 1223 K and has been used in the past for ice nucleation studies  [DeMott et al, 1999]. Neither Lamp Black 101 nor Printex 40 were post-treated. Relevant properties of these soot particles and carbon blacks are summarized in Table 3.1.  3.2.1.2  Mineral dust Two different types of mineral dust were used in our experiments: Kaolinite and  Goethite.  Kaolinite particles were purchased from Fluka Chemika (purum; natural grade).  Listed previously in Table 1.1, Chapter 1, the chemical formula for Kaolinite chemical is Al4Si40io(OH)8.  Goethite, a-FeOOH, samples were obtained from Ward Scientific and  originated in Minnesota. They were crushed and sieved prior to use.  3.2.1.3  Oi-oxidizedsoot As mentioned earlier the ice nucleating properties of Lamp Black 101 after controlled  exposed to O3 were examined. Lamp Black 101 particles were deposited on a hydrophobic glass cover slide and exposed to O3 within a flow tube in which [O3] was measured with a downstream chemical ionization mass spectrometer (CIMS). This setup is illustrated in Figure 3.2. The flow tube and C I M S instrument have been described by  [Knopf et al, 2005].  In these experiments, the pressure of N2 in the flow tube was 2 - 4 Torr and O3 was generated by photolysis of O2 at 254 nm (ultraviolet source Jelight, model #600). Ozone entered the flow-tube through a movable injector. The following ozone exposures  (poyt) were used:  1.6-10" , 7.6-10" , 1 3 . 0 1 0 , 25.1-10" , 35.6-10" , 94.8-10" atm sec. This is equivalent to exposing 3  3  3  3  3  3  the soot to 80 ppb of O3 at atmospheric pressure for 0.2, 1.1, 1.9, 3.6, 5.1, and 13.7 days, respectively. A n O3 concentration of 80 ppb at atmospheric pressure corresponds to relatively pollute atmospheric conditions  [Finlayson-Pitts and Pitts Jr., 2000].  O3 was detected as O3" in the mass spectrometer after its chemical ionization by SFe". SF " was generated by passing a trace amount of S F in about 1000 STP cn^min" N through a 1  6  6  2  Chapter 3: Experimental technique, 25 2 l 0  P o a-source ( N R D , model Po-2031).  This process can be summarized by the following  equations: N  a 2  -  )N^+e-  p a r , i d e s  Eq. 3.1  SF +e-^-SF 6  Eq.3.2  6  SF -+0 ^ S F + 0 J 6  3  Eq.3.3  6  Reaction of ozone with SF " is a pseudo-first order reaction in which the amount of SF " 6  6  remaining is defined by [SF -] =[SF -] exp(-fe[o 6  (  6  0  3>fa6  ])  Eq.3.4  where k is the reaction rate constant of (1.7±0.5)-10" cnr^molec'V [Catoire et al., 2001], tisi 10  the reaction residence time of 0.004 s, and  [Oy ] ab  1  is the quantity of ozone. Reaction residence  time was determined from the CIMS chamber volume (8.16 cm ) and the total flow rate. 3  Lens & Pin Hole  Lens""  Chemical Ionization Mass Spectrometer  Exit Lens  2(g) Carrier Movable ? Injector N  G  Quadrupole Filter  Flow Tube  8  Dynode  3E  mk  Coolant In  Pre Filter  Hydrophobic Coolant slide Out with soot  S F  6  + N  2(g) ;orr Rotary Pump 3  lO^Torr Turbo Pump  Filter  10 Torr Turbo Pump 4  Chaneltron multiplier  9  i  s  I ft  a Figure 3.2  Sketch of the flow-tube coupled to the Chemical Ionization Mass Spectrometer (not to scale). aa e  Chapter 3: Experimental technique, 27 3.2.2  Slide preparation  The bottom surface of the flow cell, which supported the particles, consisted of a glass cover slide treated with dichlorodimethylsilane (DCDMS) to make a hydrophobic layer, which reduced the probability o f ice nucleation directly on the surface.  Prior to the treatment with  D C D M S the glass slide was cleaned with a dry ice cleaning system (Sno Gun-II™, Va-Tran Systems) to remove any coarse impurities from the slide surface. Each slide was immersed in a piranha solution (3:1 mixture by volume of sulphuric acid and hydrogen peroxide) for approximately 5 minutes, rinsed in high purity water (distilled water further purified with a Millipore system Simplicity 185, 18.2 Mil) and methanol ( H P L C grade), and dried with a flow of purified N (Spiwestek, Static Prevention Inc., SF4700 HC). Each glass slide was treated with 2  a dry ice cleaning system and rinsed with high purity water and methanol for a second time. Once it was dried with the N2-ionized gas, the clean slide was placed in a glass chamber. The treatment with D C D M S involved placing the slides in an airtight chamber with 2-3 droplets of D C D M S solution (Fluka, 5% D C D M S in heptane). The slides were not in direct contact with the droplets, rather the D C D M S would coat the glass slides via vapour deposition. The resulting substrates had a contact angle o f - 1 0 0 ° , significantly better than that of a droplet on a clean glass slide (-65°). This is illustrated in Figure 3.3. The images were obtained after a droplet of water was deposited on the substrate surface. The droplet was viewed with an optical microscope from the side.  (a)  Figure 3.3 glass slide.  (b)  Image of a droplet positioned on top of a (a) clean glass slide and (b) a silanized  Chapter 3: Experimental technique, 28 3.2.3  Sample preparation  A l l samples were prepared and the flow cell constructed within a filtered air laminar flow hood.  This greatly reduced the possibility of contamination of the samples by ambient  atmospheric and laboratory particles. Soot or dust particles were deposited on a hydrophobic glass slide (the bottom surface of the flow cell) using the following technique. The dry soot or dust particulates were placed in a glass vessel immersed in an ultrasonic bath. A flow of N  2  (99.999 %) was passed through the glass vessel, and vibrations from the ultrasonic bath caused the dry particles to be suspended in the flow of N . 2  This flow was then directed at the  hydrophobic glass slide, and the soot or dust particles were deposited on the slide by impaction. Soot agglomerates or dust particles deposited on the substrate were always less than 40 pm in diameter (Figure 3.4). The optical resolution limit of the microscope was ~1 pm. A typical sample held between 200 to 800 individual particles, a majority of which were between 1 and 10 pm in diameter.  Particle size is characterized in terms of the surface mean diameter and  geometric mean diameter though out the thesis as defined by Reist [1993].  140 i  • i '» i ' . i . • i ». i. * — i — • .i ' i  6  9  12  16  18  21  24  27  30  33  36 39  Diameter/microns  Figure 3.4  Average size distribution of eight different samples of Degussa FW2 (Channel  Type Black) deposited on hydrophobic glass slides.  Chapter 3: Experimental technique, 29 3J  Flow cell setup  The flow cell was similar in design to the one used previously to measure super-micron organic and mixed organic-inorganic particles  [Pant et al., 2004; Parsons et al, 2004a; Parsons  et al, 2004b] and is illustrated in Figure 3.5. The flow cell was positioned on a cooling stage. The temperature of the cooling stage and hence the flow cell was regulated with a refrigerating circulator (Thermo Neslab ULT-95). The hydrophobic slide was positioned inside the cell body which  was  constructed  out  of  aluminium.  An  insulating  spacer,  made  from  polychlorotrifluoroethylene (PCTFE), was placed between the hydrophobic glass slide and the  (a)  1 Ox Objective Lens  Sapphire window  Viton o-rings  (b)  Figure 3.5 the  flow  (a) Three-dimensional cut away illustration of the flow cell, (b) Cross-section of cell  and  location  polychlorotrifluoroethylene).  of  microscope  objective  (Al=aluminium,  PCTFE  =  Chapter 3: Experimental technique, 30 flow cell body. This ensured that the coldest portion of the flow cell was the glass substrate (by -10 K), thus preventing unwanted ice nucleation in other parts of the cell. A l l seals within the cell were made with Viton O-rings. The upper portion of the cell body and the inlet and outlet were made from stainless steel. A sapphire window (1 mm thick), positioned at the top of the cell body, was used to monitor any changes on the bottom surface of the cell. A reflected-light microscope (Zeiss AxiotechlOO) equipped with a lOx magnifying lens was coupled to a high-resolution industrial monochrome digital video camera (Sony, XCD-X700) which captured images of the particles deposited on the hydrophobic slide during the course of a typical experiment.  3.4  Humidity control  A flow of humidified gas was introduced to one side of the cell and exited on the other where its frost point was measured with a frost point hygrometer (General Eastern). From the frost point measurements,  the water vapour pressure  (pmo) was calculated using the  parameterization of Marti and Mauersberger [1993], defined by Eq. 2.3. A flow of humidified gas was generated by passing a flow of He (99.999 %) over a reservoir of ultra-pure water (distilled water was further purified using a millipore system). The desired pmo was adjusted by altering the temperature of the water reservoir and diluting the humidified flow with a second flow of dry He.  A continuous and constant flow of between 1900 to 2100 c m min" (at 3  1  273.15 K and 1 atm) was maintained throughout the course of the experiments. The He gas used in these experiments was first passed through a trap containing molecular sieve (1/16" pellets, Type T4A) at 77 K and then through a 0.02 pm filter (Anodisc 25). The R H ; within the cell was calculated with the following equation: RH  t  =  °—100%  Eq.3.5  Ph2  Pice (Tcell)  where Pi e(T u) is the saturation vapour pressure o f ice at the temperature of the bottom surface C  ce  of the flow cell. pi {Tceii) was calculated using the parameterization of Marti and Mauersberger ce  [1993], and pmo was calculated as discussed above. The relative humidity with respect to ice corresponds to the saturation ratio such that RHi = S, -100%  Eq.3.6  Chapter 3: Experimental technique, 31 3.4.1  RHi ramp experiments In most nucleation experiments, the RHj was ramped from below 100% to water  saturation by decreasing the temperature of the cell at 0.1 K min", and maintaining a constant 1  PH20  inside the cell. Typical experimental  RHi  trajectories are illustrated in Figure 3.6 for four  different initial temperatures of 258 K, 253 K, 248 K, and 243 K. For the remainder of the document these experiments will be referred to as RHi ramp experiments. Images of the soot or clay particles were recorded digitally every 10 seconds or ~0.017 K, while simultaneously recording pmo and the cell temperature. From these images the RH; at which water droplets or ice particles first formed was determined (i.e. the onset of water or ice nucleation).  -i  Figure 3.6  r  250 255 Temperature / K  Typical experimental trajectories of RHi, where temperature was reduced at a rate  of 0.1 K min", while the water partial pressure was constant. The trajectories were calculated 1  using the saturation vapour pressure of water from Koop et al. [2000] and the saturation vapour pressure of ice from Marti and Mauersberger [1993]. The arrows show the experimental trajectory.  Chapter 3: Experimental technique, 32 3.4.2  Constant R H i experiments  Nucleation experiments with long observation times and at constant RHi were carried out in order to constrain, as much as possible, the heterogeneous nucleation rate coefficient of ice on soot in the deposition mode (see Chapter 4.2.4 for a further discussion). For the remainder of the document these experiments will be referred to as  constant RHj experiments.  In these  experiments n-hexane soot (air/fuel ratio = 2.4) was employed. This is the most oxidized soot and therefore will have the greatest ice nucleating potential.  A s mentioned above, the  International Steering Committee for Black Carbon Reference Materials has recommended using n-hexane soot as a model for soot in the atmosphere.  In the  constant RHj experiments, the  temperature of the particles was held at -248 K , while the relative humidity was held at 124 ± 4 % RHi, which is just below water saturation. The particles were held at these conditions for an extended period of time (approximately 8 hours) and were monitored to determine i f ice nucleated during this long observation time.  3.5  Cell temperature & temperature calibration  A Pt-100 resistance temperature detector (RTD) from Omega was embedded within the aluminium base to measure the temperature of the bottom surface o f the cell. The R T D was calibrated against the dew point or ice frost point within the cell, as done previously  [Middlebrook et al., 1993; Parsons et al, 2004b]. In order to determine the difference between the measured R T D reading and the temperature of the bottom surface of the cell, the temperature of water droplets or ice particles was varied while images of the particles were simultaneously recorded. From these images the temperature at which the size of the droplets or ice particles did not change was determined.  While their size remained constant, the liquid droplets or ice  particles were in equilibrium with the water vapour inside the cell, which was precisely known. Hence it was possible to determine the offset temperature between the measured R T D reading and the temperature of the liquid droplets or ice particles formed on the bottom surface of the cell. As an example, Figure 3.7 illustrates the change in the surface area of liquid droplets as a function of RHi, frost point, R T D readings, and the time. First, the temperature of the cell was decreased (A) at a rate of 0.1 K min" until water droplets were formed (B) and grew (C). Then 1  Chapter 3: Experimental technique, 33 the cell temperature was slowly ramped up (rate = 0.1 K min" ) until the water droplets were 1  observed to shrink (D). Next the cell temperature was slowly decreased again until the water droplets started to grow. From these observations the temperature at which water droplets were in equilibrium with the gas-phase water vapour was determined. In Figure 3.7 water droplets were observed to nucleate at ~ 10:41 am, 254.59 ± 0.03 K (RTD), and a frost point of 256.3 ± 0.2 K . During the calibration process, the surface area of the liquid droplets remained constant a total of three times (shaded portions of the plots in Figure 3.7 between -10:53 to 10:58 am, 11:02 to 11:08 am, and 11:16 to 11:25 am). These regions were assigned by considering the scatter in the data for which the average R T D reading was determined.  This approach was compared against a more quantitative technique in which a  polynomial function was fitted over the surface area data. The derivative provided R T D values well within uncertainty. This indicates that the visual assignment of the point at which the liquid droplets are in equilibrium with water vapour pressure is valid. At some point during each of these intervals the temperature of the bottom surface of the cell was equal to the dew point of the vapour, which was determined independently from the hygrometer measurements. Although the hygrometer measurements gave the ice frost point, this reading was converted into a dew point using the saturation vapour pressure of water from Koop  et al. [2000] and the saturation vapour pressure of ice from Marti and Mauersberger [1993]. A n average R T D offset of 0.29 K was determined from all three time intervals.  B y applying this  offset correction to the R T D reading at the onset of liquid droplet formation, the temperature of the bottom surface of the cell was determined to be 254.3 ± 0.3 K . The uncertainty was determined by incorporating hygrometer error (± 0.2 K ) with the spread in the calibration readings (based on the shaded portions of the plot). The relative humidity for the formation of the liquid droplets at water saturation was determined based on the correct temperature reading and the frost point.  The uncertainty in onset RHj was determined by incorporating the  hygrometer error and calibration error to give the RHj of 121.6 ± 4.7% at the onset of liquid droplet formation. This calibration process over liquid water saturation is illustrated in Figure 3.8 through a sequence of images. Each image ties in with the above calibration description (i.e. image (B), taken at 10:40:49 am, illustrates the initial condensation of water droplets on a substrate coated with Degussa FW2).  Chapter 3: Experimental technique, 34 The multi-step temperature calibration for experiments in which ice crystals were observed to form is very similar to the above technique. In the event of deposition nucleation of ice, the R T D was calibrated against the ice frost point within the cell rather than against liquid water saturation.  Chapter 3: Experimental technique, 35  c  2  o E ca CL) CO CD  t  CL  2 Q T  10:30  '  I  10:40  10:50 1  256.5  11:00  255.5  2  255.0  Cl) Q_  E  11:20  '  w%  1  11:30  -  Frost Point.  V////.  256.0  2  11:10  lip  v/y/yy// ////////  vyy/// vy/V  254.5 254.0  Corrected Cell. Temperatire -  IP  253.5 253.0  / / / /  m  130  125-1 120  11:30  10:30  Figure 3.7  This figure illustrates the change in surface area of liquid droplets condensed on a  hydrophobic slide in relation to RHi, dew point, R T D , and cell temperature (corrected for offset). As the cell temperature was ramped down (A) at a constant rate o f 0.1 K min" , the R H i over 1  deposited particles increased. Liquid droplets were first condensed when R H = 100% (B). They W  continued to grow in size (C) until R H < 100% (D). Once again the cell temperature was W  decreased at a constant rate and the droplet surface area increased above 100% R H . W  Chapter 3: Experimental technique, 36 A) 10:35:08 am r u = 254.7 K , R H i = 118%  C) 10:47:58 am 7/ = 254.0 K , RHj= 125%  B) 10:40:49 am r = 254.3 K , RHj= 122%  D) 11:01:03 am r n= 254.6 K , RHj = 119%  ce  c e n  Figure 3.8  ceI1  ce  Images were recorded as the temperature of the cell was ramped down and up at  constant p o- Labels (A) to (D) correspond to labels in Figure 3.7. The black deposits are H2  Degussa FW2 soot on a hydrophobic surface. Each image is accompanied by experimental time, r n, and the RHj. ce  Chapter 4: Soot, 37 Chapter 4  Ice nucleating properties of soot  4.1  Introduction  Although soot is abundant in the atmosphere  [Cooke and WUson, 1996; IPCC, 2001], its  contribution to the formation of ice under lower tropospheric conditions is yet to be quantified and remains uncertain. This section details the results obtained through  RHi ramp experiments  for Lamp Black 101, Degussa FW2, Printex 20, and n-hexane soots in the temperature range of 243 to 258 K and the constant RH,  experiment with n-hexane soot. Data collected on nucleating  properties of Lamp Black 101 following controlled oxidation with ozone is also presented. The contribution of soot to formation of ice in the lower troposphere is quantified by evaluating the upper limit to the nucleation rate coefficient,  J" . The lower limit to the contact angle for p  h el  formation of a critical-size ice germ on soot is then determined by applying heterogeneous nucleation theory (Chapter 2).  J"  p h el  and the contact angle are used to determine the maximum  number of ice particles that can nucleate on soot particles in the atmosphere at 248 K in the deposition mode. This value is compared against field data in order to determine whether the ice nucleating contribution of soot can account for the presence of the ice phase at lower tropospheric conditions.  4.2  Results  4.2.1  R H j ramp experiments  In a typical experiment, pmo was held constant while the temperature of the cell was reduced in order to increase the RHi within the flow cell. The temperature was decreased until either water droplets or ice particles were observed. The RHi at which either water droplets or ice particles were observed is illustrated in Figure 4.1 for the blank hydrophobic glass slide as well as the different soot samples. The dashed lines in the figure represent water saturation (i.e. relative humidity with respect to water is 100%). The open symbols indicate that water droplets  (B)  (A) I  14.0  (O  I  130 120 110 4  - - Water saturation !> Blank onset  • o H  '100 (D)  A N-hexane (diffusion flame) onset N-hexane (aff=0.53) onset 1 1 > + ' 1' 1(E)  •  <1 N-hexane (a/f=2.4) onset 1 1 + 1 1 ' (F)  140 130 120. 1 TO • 100  240  Figure 4.1  O  Lamp Black 101 onset 245  250  255  260  240  •  o  Printex 40 onset 245  250 255 Temperature/ K  260  240  Degussa FW2 onset 245  250  255  260  The RHj at which liquid water droplets or ice particles were first observed as the RHj inside the cell was slowly increased.  3  The open symbols indicate that water droplets were first observed (sometimes ice was observed to form after water droplets condensed) and the solid symbols indicate that only ice particles were observed with no indication o f the formation of water droplets prior to ice formation. Panel (A) illustrates the results from a control experiment, where no particles were deposited on the hydrophobic substrate. Panel (B) to (F) illustrates the results from the six soot types listed in Table 3.1.  Oo  Chapter 4: Soot, 39 were first observed and the solid symbols indicate that only ice particles were observed with no indication of the formation of water droplets prior to ice formation.  From this information  conclusions are made on the ice nucleating properties of various soots in the deposition mode below water saturation (see below). In the experiments where water droplets were first observed (open symbols), ice nucleation would occasionally occur at a later time, presumably by immersion freezing. This occurred with both the blank as well as with soot particles. However, it is not possible to determine i f the formation of ice after the formation of liquid droplets was due to the soot particles or the substrate.  Hence this information is not included. The current experimental  configuration is not well suited for investigations of immersion freezing. The results in Figure 4.1 show that at 248 K and above, water droplets, rather than ice, always appeared first in our experiments. This occurred at water saturation, as expected. From this observation it may be concluded that ice nucleation never occurred at temperatures above 248 K and below water saturation for the current experimental conditions (observation time and soot particle concentrations). If ice nucleation did occur, ice particles would rapidly grow and prevent the formation of water droplets or new ice crystals at water saturation by depleting the water vapour. A t ~ 243 K , occasionally only ice particles formed with no indication of the formation of water droplets prior to ice formation (a total of three times). However, in the few experiments where ice did form, the RHj was close to water saturation when ice nucleation was observed, suggesting water nucleation may have occurred first, followed by ice nucleation during the condensation process. In other words, for the few experiments where ice did form condensation freezing cannot be ruled out. In fact, at temperatures between 243 and 258 K all the results (including when ice nucleated first) clustered around water saturation, suggesting water saturation is a prerequisite for both water and ice nucleation.  4.2.2  Ice nucleation of soot particles oxidized by O 3  Lamp Black 101 was exposed to various amounts of ozone and then  RHj ramp  experiments were performed to test the ice nucleating properties of these soots. The results are shown in Figure 4.2. The open symbols indicate that water droplets were observed first in all experiments. The results for Lamp Black 101 exposed to ozone are the same as the results from  Chapter 4: Soot, 40 unexposed Lamp Black 101. In all cases, water droplets were first observed, indicating that ice did not nucleate below water saturation. Even after an O 3 exposure of 9.510" atm sec, which is 2  equivalent to an exposure of 80 ppb at atmospheric pressure Opolluted conditions) for 13.7 days, the results were not different from results of unexposed Lamp Black 101. Either O 3 did not oxidize Lamp Black 101 significantly or the oxidation process did not change the ice nucleating ability significantly. Further research is needed to determine the extent of oxidation of Lamp Black 101 by O 3 .  Also more research is needed to determine i f exposure to atmospherically  relevant concentrations of ozone, as well as other atmospheric oxidants such as O H and N 0  3  radicals, can modify the I N properties of other types of soot in the deposition mode as well as other modes of ice nucleation. These initial experiments show that exposure to atmosphericallyrelevant concentrations of ozone did not modify the ice nucleating ability of Lamp Black 101 in the deposition mode below water saturation.  1404  '130.4  -120: Waterline Lamp Black unexposed o Lamp Black exposed for 0.2 days A Lamp Black exposed for 1.1 days O Lamp Black exposed for 1.9 days o Lamp Black exposed for 3.6 days <3 Lamp Black exposed for 5.1 days ;>. Lamp Black exposed for 13.7 days •  110  100 240  Figure 4.2  245  250 Temperature / K  255  260  The R H i at which water droplets were observed for Lamp Black 101 soot samples  treated with a range of O 3 exposures. In all experiments water droplets were observed first (i.e. ice particles did not form unless water droplets first condensed). The exposure times indicated in the plot are equivalent to an atmospheric mixing ration of 80 ppb (see text for details).  Chapter 4: Soot, 41 4.2.3  J £ for RHi ramp experiments u  t  For the experiments where water droplets first formed, an upper limit to the deposition nucleation rate coefficient of ice on soot particles was estimated. This rate coefficient provides a quantitative measure of the ice nucleating ability, which may be used in modeling studies of ice formation in the atmosphere. Based on Poisson statistics, i f ice nucleation did not occur during the course of an experiment, an upper limit to the heterogeneous nucleation rate coefficient, J h et, u p  can be calculated with the following equation  [Biermann et al., 1996; Koop et al, 1995; Koop et  al, 1997]: r  p  1  =—In  Eq. 4.1  1-  where r is the observation time, A  s  is the total surface area available for heterogeneous  nucleation, and x is the confidence level (95 % was used). During  RHj  ramp experiments x was  approximately 60 seconds (the cell temperature changed by 0.1 every 60 seconds) and^4 ranged 5 5 2 s  from M(r to 4-10 pm . A was calculated by assuming a spherical geometry with surface area J  s  equal to 4^r , where ro is the radius of the soot particles. This is a conservative estimate as the 2  0  surface area exposed to the gas phase is in most cases larger than the geometric surface area. For instance, in the case of n-hexane (air/fuel = 2.4) the B E T surface area was found to be 156 ± 11 m -g"', which corresponds to a scaling factor of 490% larger than the geometric surface area 2  (assuming the particle was 10 um in diameter and had a density of 1.86 g-cm" ). However, this 3  area measurement may not be equal to that available for ice nucleation and is not applied to the calculation of 50 cm" sec" . 2  1  J . Based on a conservative surface area of 1 T 0 pm , J u p  h el  5  2  up el  was calculated to be  In the subsequent section the upper limit o f ice nucleation will be further  constrained in the constant RHj experiments. 4.2.4  Constant R H i experiments  During the constant RHj experiment, n-hexane soot particulates (air/fuel = 2.4) were held at -247.5 K and close to liquid water saturation for approximately 8 hours. The surface distribution of soot particles is shown in Figure 4.3. The surface mean diameter of particles in this sample was 19 pm, whereas the geometric mean diameter was 7.1 pm. The experimental  Chapter 4: Soot, 42 conditions are plotted in Figure 4.4.  The plot illustrates relative humidity fluctuations as a  function of the cell temperature and the frost point drift during the eight-hour period. During this experiment soot particles were held at 124 ± 4 % RHj, 1 to 8% below liquid water saturation.  g. o |  403020-  o| 0  WUU WW PJ P" 5  10  15  20  25  . 30  35  40  45  50  Diameter / microns  Figure 4.3  (a) N-hexane (air/fuel = 2.4) soot particles on a hydrophobic slide. The surface  area of the soot particles is 1.1-10 pm . (b) Sample size distribution during the constant RHj s  2  experiment. The geometric mean diameter is 7.1 pm, with standard deviation of 1.2 pm.  Chapter 4: Soot, 43 1200 p—-—•  13:30 1  15:00 '—•—i  16:30 1  '  18:00 1  '  19:30 i  '  21:00 r  1  Liquid water saturation .  247  J  •h  12:00  1  1  13:30  <  1  15:00  i  1  16:30  1  1  18:00  •  1  19:30  1  r*  21:00  Time  Figure 4.4  Change in relative humidity with respect to ice (black) during the constant RH  t  experiment. The liquid water saturation in terms of R H ; is also plotted (pink); it shows that water saturation was never reached. The drift in the frost point (green) and the cell temperature (blue) is also illustrated. For a period of approximately 8 hours cell temperature was held at 247.5 K and RHj = 124 ± 4%, close to water saturation. Water droplets or ice crystals were not observed to form.  Chapter 4: Soot, 44 5  2  In this experiment the surface area of soot exposed to the vapour was 1.1 TO pm . Even during this long observation time at humidities close to water saturation, no ice was observed. From this an upper limit to the heterogeneous nucleation rate coefficient was calculated using Eq. 4.1. In this case the upper limit XoJ  u p h et  was determined to be 0.1 cm" sec" . The upper limit 2  is much smaller than the upper limit calculated from the  1  RHi ramp experiments, of 50 cm" sec" , 2  1  since the observation time was much longer (8 hours compared with 1 minute). This estimate provides a better constraint on the rate coefficient of ice nucleation on n-hexane soot in the deposition mode. 4.3  Comparison with other results  There have been several measurements  of the ice nucleation ability of soot at  temperatures above 238 K . In addition there have been a few studies at lower temperatures. DeMott et al. [1999] investigated ice nucleation on Lamp Black 101 at temperatures ranging from 233 K to 213 K using a continuous flow diffusion chamber. In Figure 4.5, these results are compared with the current data. A t approximately 230 K , DeMott  et al. [1999] observed ice  nucleation only at water saturation. Hence, at warmer temperatures it is highly unlikely that they would observe ice nucleation below water saturation, which is consistent with findings presented in this thesis.  Mohler et al. [2005a] investigated ice nucleation on spark generated soot at temperatures less than 240 K using a low-temperature aerosol and cloud chamber (Aerosol Interactions and Dynamics in the Atmosphere, A I D A ) . A t temperatures between 235 K and 240 K , ice nucleation only occurred on uncoated soot particles after approaching water saturation.  The authors  commented that ice seems to form immediately in this temperature range after liquid activation of the soot particles either by condensation freezing or homogeneous freezing of the growing liquid water layer. This finding is consistent with the current studies. More recently, Mohler  et  al. [2005b] used the A I D A chamber to investigate ice nucleation at low temperatures on soot produced from combustion of propane with various elemental carbon to organic carbon ratios. If these results are extrapolated to warmer temperatures they are also consistent with findings presented in this thesis.  Chapter 4: Soot, 45 v  160  \  \  :  150  140-1  130  120  110  - - Water saturation • i Unexposed Lamp Black • DeMott et a/. ;  100 210  Figure 4.5  ;  220  230 : 240 Temperature / K  250  260  A comparison of current results for Lamp Black 101 (open squares) with those of  DeMott et al. [1999] (filled triangles). Our data points correspond to the conditions at which water droplets were observed using soot particles ranging in size from 1 to 40 um in diameter. In these experiments water droplets were always observed first. If ice did form it was only after the appearance of water droplets. The results from  DeMott et al. correspond to the onset for which  1% of Lamp Black soot particles (a number mean diameter of 240 nm) nucleated ice.  DeMott [1990] investigated ice nucleation on soot produced by combustion of acetylene at temperatures between 233 K and 253 K using an expansion cloud chamber. Due to the experimental design and experimental conditions, mainly condensation and immersion freezing were investigated. DeMott [1990] commented that there was some evidence of ice formation by deposition, but ice certainly did not precede cloud droplet formation significantly. One of the conclusions from this study was that immersion freezing nucleation is an efficient ice nucleation process after water has condensed on soot particles.  Chapter 4: Soot, 46 Diehl and Mitra [1998] investigated ice nucleation on soot produced by the combustion of kerosene. Combined deposition and condensation freezing (deposition/condensation freezing) were studied in a single experiment using a soap film method. Ice nucleation occurred in these experiments at temperatures as high as 253 K . In these experiments, the relative humidity was not  measured,  so  a  direct  comparison with  present  results  is  difficult.  A l l the  deposition/condensation experiments may have been carried out slightly above water saturation and condensation freezing may have dominated. The authors also studied immersion freezing and contact freezing, and they found that kerosene-bumer exhaust particles are effective ice nuclei in these freezing modes. Finally, Gorbunov et al. [2001] investigated soot produced by the combustion of benzene and toluene, as well as soot produced by thermal decomposition of benzene using a cloud chamber at temperatures ranging from 253 to 268 K . A l l experiments were carried out close to liquid water saturation: saturation with respect to liquid water was equal to 1.02 ± 0.02. It was found that the fraction of aerosol particles forming ice crystals was influenced by the concentration of surface chemical groups that can form hydrogen bonds with water molecules. A large difference in the ice-forming activity (3 orders of magnitude in the fraction of soot particles forming ice crystals) was observed for soot aerosols obtained with different generators. Soot particles produced by combustion of benzene and toluene were very potent ice nuclei, whereas soot particles produced by thermal decomposition of benzene were poor ice nuclei. They concluded that highly oxidized soot particles are extremely efficient ice nuclei.  The  difference between the present results and the results from Gorbunov et al. [2001] may be due to a difference in experimental conditions: current research focuses on ice nucleation below water saturation whereas the work presented by Gorbunov et al. [2001] was carried out at or slightly above water saturation. Alternatively, the soot particles studied by Gorbunov et al. [2001] were more effective I N than the soot investigated in this study. Combining all the previous results and those presented in this thesis, it appears at temperatures above 243 K and below water saturation, ice nucleation on many types of soot particles is not efficient  [DeMott, 1990; DeMott et al, 1999; Mohler et al, 2005a; Mohler et al,  2005b]. In contrast, once the RHj is above liquid water saturation, water can condense on soot particles, and then most types of soot may be important ice nuclei in the condensation or immersion mode  [DeMott, 1990; DeMott et al, 1999; Diehl and Mitra, 1998; Gorbunov et al,  2001; Mohler et al, 2005a; Mohler et al, 2005b].  Chapter 4: Soot, 47 4.4  Contact angle  Although other authors have utilized experimental data to determine the contact angle of an ice embryo in the immersion mode on various insoluble I N [Hung et al, 2000; Archuleta et ah, 2005], values pertaining specifically to the formation of ice on soot particles in the deposition mode are yet to be published. The contact angle, 9, is thought to be independent of RHj, hence 6 obtained from different insoluble species, that nucleate ice at different RHj, may be compared. Here, heterogeneous nucleation theory (Chapter 2) is employed to determine the lower limit to the contact angle of the ice germ formed on soot under conditions defined by the constant and  ramp RHi experiments. The experimentally-determined J"  P h t  is used to constrain the contact angle of the critical-  size ice embryo formed on the surface of soot. The lower limit to the contact angle is determined by first finding the limiting value to the activation energy, defined as f  A G , =-kTln  TUP  °hel  V  Eq. 4.2  A  where cell temperature, T, is 247.5K, J"h el is 0.1 cm" sec"' , and A is the pre-exponential term p  2  {10 } cm" sec"' [Young, 1993]. These conditions correspond to the constant RHj experiment (at 26  2  which Si = 1.24). Based on these definitions AG* , for formation of a single ice embryo is C  2.1 TO" J. If a single critical-size spherical-cap ice germ forms on an insoluble soot surface, 19  independent of lattice strain (i.e. e = 0), the size of the germ is calculated with Eq. 2.20. The lower limit radius of critical size is 9.4 nm. The lower limit to the contact angle may be calculated by re-evaluating Eq. 2.22 for the geometric t e r m , X ^ ) , such that .... 3 [m^.lnS,.] . /("») = — " ' AG , 2  3  16  Eq. 4.3  ac  KG  UV  In the above equation TV, is the molecular concentration of water molecules within the ice germ and equals to 3.1 T O  28  molecules r n . This value is calculated based on ice density of 0.92 g cm" 3  3  and is a reasonable estimate, as the density of hexagonal ice varies only slightly with temperature [CRC, 2001-2002].  In order to estimate the lower limit to the contact angle, the interfacial  tension between the ice and vapour phase of water was assumed to be 106.5 mJ m" [Kay et al., 2  2000; Pruppacher and Klett, 1997]. It is of importance to note that cr is yet to be determined iv  Chapter 4: Soot, 48 experimentally. For this reason, a should be considered highly uncertain [Kay et al, 2000]. If n  the ice germ forms on a flat insoluble soot surface, the geometric term, f\ m), is defined by the following parameterization m -3m +2 3  f(m) =  Eq.4.4 4  where m = cos 6. With these assumptions, the lower limit to the contact angle for the formation of an ice germ on the flat soot surface is 24°. The inclusion of the lattice misfit parameter in the above calculation is expected to alter the contact angle.  However, large variability in soot  morphology makes the misfit parameter difficult to estimate - hence, it is neglected here. This approach was also applied to several  ramp RHi experiments for which the total soot  surface area was determined. Table 4.1 summarizes parameters involved in the calculation of the lower limit to the contact angle (i.e. onset temperature, onset RHj, and limiting values of J , het  L\G* , and r). acl  The calculated lower limit of the contact angle is plotted against temperature in  Figure 4.6. The data reflects the dependence of the contact angle on the liquid saturation ratio. As the onset temperature decreases, the lower limit to the contact angle for formation of the ice embryo increases with liquid water saturation.  Chapter 4: Soot, 49 Table 4.1  Onset parameters pertaining to formation of a single ice crystal in multiple ramp  RHi experiments for which Jf  el  was determined.  The confidence level for J"£, is 95%. The  limiting values of A G * , , r , and the contact angle are also included. A G * C  the assumption that the pre-exponential term, A, equals to 10  26  C(  was evaluated under  cm^sec" . The size of the critical1  size ice germ was evaluated based on Eq. 2.20 pertaining to the maximum Gibbs free energy of formation. Finally, a lower limit to the contact angle was determined using Eq. 4.4.  Soot type  Printex 40  Lamp Black  Lamp Black-03  n-hexane (0,X)  n-hexane (6,X)  n-hexane (11,X)  Degussa F W 2  Onset Temperature  Onset RH,  EKiilfB IllPil 258.5±0.2 252.9±0.2 248.4±0.2 242.2±0.2 257.8±0.2 253.1±0.2 249.0±0.2 243.1±0.2 248.8±0.2 248.5±0.2 248.4±0.2 248.2±0.2 248.1±0.2 248.1±0.2 247.9±0.2 243.8±0.2 243.8±0.2 243.3±0.2 243.2±0.2 243.1±0.2 257.7±0.2 253.2±0.3 248.2±0.2 243.3±0.2 258.0±0.2 253.0±0.2 247.7±0.3 243.5±0.3 256.1±0.2 250.5±0.2 248.5±0.3 243.8±0.2 257.6±0.2 253.5±0.3 249.3±0.2 243.5±0.2  113±3 123±4 127±4 138±4 117±3 121±4 124±3 133±4 124±3 128±3 129±3 127±2 129±3 128±3 131±3 131±4 131±3 133±3 136±4 135±4 116±4 121±4 127±3 135±4 119±3 123±3 129±5 135±5 119±3 125±3 127±5 132±3 114±3 122±6 127±3 133±3  • /*i||pi*T- i'« % i *i£ £ : ;  :  :  :  :  Contact \ngle  Tip het  AG; ,  (cm' sec"')  (j-io'- )  (nm)  IMS!  53 23 23 131 52 52 52 52 68 68 122 58 84 62 5 68 68 62 84 58 17 17 17 17 13 13 13 13 18 18 18 140 100 28 19 25  1.99 1.98 1.95 1.84 1.99 1.95 1.92 1.88 1.91 1.91 1.89 1.91 1.90 1.90 1.99 1.87 1.87 1.86 1.87 1.43 2.04 2.00 1.96 1.92 2.04 2.00 1.96 1.92 2.01 1.97 1.95 1.85 1.97 1.98 1.96 1.91  16 10 9 6 13 10 10 7 9 8 8 9 8 8 8 8 8 7 7 7 13 10 8 7 11 9 8 7 11 9 9 7 15 10 9 7  18 23 25 28 20 22 23 27 24 25 25 25 25 25 27 26 26 27 27 27 20 23 25 27 22 24 26 27 22 24 25 26 19 23 25 27  J  :  |lJ  Chapter 4: Soot, 50 30  1  1  •  1  1  •  1  1  ... ——  i  •  1  r  '  i  1  n  X  28V)  0)  a  Ui  26  %* x  0) "D  .2 24  :•  C  -  X  X  $  ro "G iS 22 c o o  X:  *  *  a)  £ 20 o< * -—  E  X  1:84  I 164  14 240  x  Spot - ramp RH.experiments  n  Soot - eonsf. RH experiment  242  244  246  248  250  252  254;  X  258  256-  -  260:  Temperature / K Figure 4.6  Lower limit to the contact angle vs. temperature.  Here a critical-size, spherical-  cap ice germ was assumed to form on an insoluble soot surface approximated to be flat. The contact angle was determined for N = 3.1 T O  28  {  m" and o i = 106.5 mJ m" using Eq. 4.3 and Eq. 3  2  v  4.4. Plotted data corresponds to the constant and ramp R H i experiments for which the soot surface area was known. Since the contact angle is a limiting value, it depends on liquid water saturation, and increases at lower temperatures.  4.5  Atmospheric implications  The maximum number of ice particles that can be produced in the atmosphere at 247.5 K and at RHj = 124% can be estimated from the nucleation rate coefficient, place the magnitude of  J"  P H ET  J£. U  T  The purpose is to  determined experimentally into an atmospherically-relevant context.  Since soot quantities vary with season, source proximity, and altitude, the maximum possible  Chapter 4: Soot, 51 number of ice particles depends on carbon mass concentrations  dictated by specific  environmental conditions. The subsequent treatment focuses on conditions resembling a typical US urban-influenced rural area [Seinfeld andPandis, 1998]. The following equation can be used to estimate the maximum number of ice particles that can be produced during a specified time period [Pruppacher andKlett, 1997],  n = n [l - exp(- J%A T)] ice  soot  Eq. 4.5  P  where nice is the number density of ice particles produced (litre" ), nsoot is the number density of 1  soot (litre" ), A is the surface area of a single soot particle (cm ), and r is the total time 1  2  p  (seconds). For ice nucleation in the atmosphere it was assumed rwas approximately 60 minutes, n t soo  was 1.5T0 litre" , and Av was 1.3T0" cm . A value of 1.5T0 litre" for nsoot was calculated 5  1  9  2  5  1  by assuming a soot radius of 0.1 pm, a soot density of 2 g cm" , a geometric surface area for soot, 3  and an elemental carbon mass concentration in the atmosphere of approximately 1.3 TO" g m" 6  3  (which corresponds to urban-influenced-rural areas [Seinfeld and Pandis, 1998; Shah et al, 1986]). A value of 1.3TO" c m for A was calculated based on a geometric surface area and a 9  2  p  soot radius of 0.1 pm.  Soot radius was obtained from Blake and Kato [1995]. This is in  excellent agreement with more recent work [Berner et al, 1996; Longfellow and Ravishankara,  2000]. With these assumptions, a maximum number density of ice of 0.07 litre" was obtained 1  from Equation 4.5.  This approach can be applied to different carbon mass concentrations.  Illustrated in Table 4.2, the scenarios include conditions characteristic of the upper troposphere, free troposphere, US remote areas, and polluted urban areas. Based on field observations, Meyers et al. [1992] developed an empirical relationship between the number concentration of ice nuclei and ice supersaturation N = exp {a + b[\00(5,. -1)]}  Eq. 4.6  id  where Nicj (litre" ) is the number of ice crystals predicted as a result of deposition nucleation or 1  condensation freezing, a = -0.639, and b = 0.1296. This parameterization predicts that at 124 % RHi, the number of ice nuclei in the atmosphere is ~ 12 litre" . Ice nucleation on soot particles 1  below water saturation (with properties similar to the soot studied in our experiments and n , = soo  1.5T0 litre" ) cannot account for this number density at 124 % RHj and 247.5 K . This suggests 5  1  atmospheric aerosols other than soot contribute to the number density of ice crystals observed in the field.  Table 4.2  A summary of the number density of predicted ice particles per litre nucleated by soot particles for five locations: upper  troposphere, free troposphere, U S A remote area, U S A urban-influenced rural area, and U S A polluted urban area. Elemental carbon mass concentrations vary with location and latitude.  [Blake and Kato, 1995]. The  Soot radius was assumed to be 0.1 pm  experimentally-determined upper limit to the heterogeneous ice nucleation rate coefficient was 0.1 cm" sec  at 124% RHj and 247.5 K .  2  The maximum amount of ice particles that can be produced for each scenario increases with soot density. Under these conditions  Meyers et al. [1992] predicts that at 124% RHj the number of ice particles in the atmosphere is approximately 12 litre" . 1  Elemental carbon Scenerio  mass concentration  Upper limit to the Number density of soot per litre (litre" )  Reference  number density of ice  1  (gm"Y Upper Troposphere  1.8xl0"  Free Troposphere  25xl0"  U S A remote area  0.5xl0"  9 a  9 6  per litre (litre' ) 1  215  2.9xl0  3  8.9xl0"  5  1.2xl0"  3  4  1.3xl0" ^  1.55xl0  5  0.065  3.8xl0"  4.54x10  s  0.19  c  Hauglustaine [1996]  b  0.025  5.97xl0  6  "Blake and Kato [1995]  'Seinfeld and Pandis [1998]  Seinfeld and Pandis, [1998]; Shah [1986]  d  U S A urban-influenced rural area U S A polluted urban area  6  6 e  e  Seinfeld and Pandis [1998]  Chapter 4: Soot, 53 4.6  Conclusions  The results presented in this section demonstrate that the soot types investigated here are poor deposition-mode ice nuclei between 243 to 258 K , below water saturation. In the present study soot particles were exposed to conditions resembling those of the lower troposphere. A t 248 K and above, water droplets always nucleated first. Consequently, it was concluded that at these temperatures ice nucleation does not occur below water saturation for the defined experimental conditions. A t ~ 243 K , ice formed first occasionally with no indication of the formation of water droplets prior to ice nucleation. However, even at these temperatures the RHj was close to water saturation when ice nucleation was observed. The results of ice nucleation on Lamp Black 101 exposed to ozone were similar to the results from unexposed Lamp Black 101. Even after an O3 exposure of 9.5T0" atm sec, which 2  is equivalent to an exposure of 80 ppb at atmospheric pressure for 13.7 days, the results were not significantly different from results of unexposed Lamp Black 101. A n experiment was carried out at a constant RHj and over a long observation time (8 hours) on n-hexane soot (air/fuel = 2.4). Even after a long observation time at 248 K and close to water saturation (RHj = 124 ± 4%), no ice was observed. From this measurement an upper limit to the heterogeneous nucleation rate coefficient of 0.1 cm" sec" was calculated. 2  1  Combining all the previous results and the current data, it appears that below water saturation at temperatures above 243 K , ice nucleation on many types of soot particles is not efficient  [DeMott, 1990; DeMott et al., 1999; Mohler et al, 2005a; Mohler et al, 2005b]. In  contrast, once the RHj is above liquid water saturation, water can condense on soot particles, and then most types of soot may be important ice nuclei in the condensation or immersion mode  [DeMott, 1990; DeMott et al., 1999; Diehl and Mitra, 1998; Gorbunov et al, 2001; Mohler et al, 2005a; Mohler et al, 2005b]. In order to further quantify the ice nucleating abilities, heterogeneous nucleation theory was employed to determine the contact angle of an ice germ formed under conditions defined by  constant and ramp RHj experiments.  In the  constant RHj experiment, the upper limit to  heterogeneous nucleation rate was employed to determine the lower limit to the contact angle of 24°. The lower limit to the contact angle for RHi  ramp experiments was also determined. Since  the J% is limited to liquid water saturation, the contact angle increased at lower temperatures. t  Chapter 4: Soot, 54 The consequences of the ice nucleating abilities of soot in the lower troposphere were considered.  Based on results obtained from the  constant RHj experiment, J  u p h el  was used to  determine the upper limit number density of ice particles under typical atmospheric conditions. For conditions resembling the U S A urban-influenced rural areas, the number of ice particles produced by the characteristic number density of soot of 1.5T0 litre" was at most 0.07 particles 5  1  litre" . This result corresponds to the number of ice particles formed as a result of deposition 1  nucleation below liquid water saturation. The value was found to be significantly smaller than the number of ice particles predicted at RHj = 124% by  Meyers et al. [1992], and implied that  soot will not significantly alter the formation of clouds in the lower troposphere at temperatures above 243K, below water saturation.  Chapter 5: Mineral Dust, 55 Chapter 5  Ice nucleating properties of mineral dust: Kaolinite and Goethite  5.1  Introduction  The deposition-mode ice nucleating properties of Kaolinite and Goethite particles were investigated between 239 to 258 K , below liquid water saturation.  Kaolinite particles are  believed to be a significant component of dust particles in the atmosphere [Glaccum and Prospero, 1980; Pye, 1987], consisting of up to 25% of the total clay concentration in Tamanrassett, Sessali, and In Guezzam, Africa  [Goudie and Middleton, 2001]. Goethite is one  of the most stable iron oxides found in soils at ambient temperature [Cornell and Schwertmann, 2003]. The iron content of soil dust varies globally, but on average makes up 3%  [Jickells et al.,  2005; Usher et al, 2003]. This section details the results obtained from  RHj ramp experiments for Kaolinite and  Goethite in the temperature range of 239 to 258 K . The onset of ice formation is quantified with Jhet-  The contact angle for the formation of a single ice embryo on Kaolinite and Goethite is also  determined for each experiment. B y utilizing the contact angle, the atmospheric implications of these mineral dust results are considered.  Chapter 5: Mineral Dust, 56 5.2  Results  5.2.1  Ice nucleation on Kaolinite  During a typical  RH ramp experiment, pmo was held constant while the temperature of t  the cell was reduced in order to increase the RHj within the flow cell. The temperature was decreased until either water droplets or ice particles were observed. The RHj at which either water droplets or ice particles were observed is illustrated in Figure 5.1 for the blank hydrophobic glass slide as well as for Kaolinite. The dashed line in the figure represents water saturation (i.e. relative humidity with respect to water is 100 %). The open symbols indicate that water droplets were first observed and the solid symbols indicate that only ice particles were observed with no indication of the formation of water droplets prior to ice formation. From this information conclusions are drawn on the ice nucleating ability of Kaolinite in the deposition mode (only) below water saturation. In the experiments involving Kaolinite particles at temperatures above ~ 252 K , water droplets were first observed (open symbols), with ice nucleation occurring at a later time, presumably by immersion freezing. However, whether or not the formation of ice following the formation of liquid droplets was due to the Kaolinite particles or the substrate cannot be determined. Consequently, this information is not included. However, the results in Figure 5.1 show that at temperatures above ~ 252 K , water droplets, rather than ice, always appeared first in  RHi ramp experiments involving Kaolinite. This occurred at water saturation, as expected. This result implies that for the present experimental conditions ice nucleation never occurred at temperatures above ~ 252K and below water saturation. At temperatures below ~ 252 K ice nucleation was consistently observed. Ice particles grew rapidly and prevented the formation of water droplets and new ice particles by depleting the water vapour. Images illustrating formation of ice on Kaolinite are shown in Figure 5.2. These ice particles nucleated with no indication of the formation of water droplets prior to ice formation. A s is shown in Figure 5.1, ice formed at RHj below water saturation. Additionally, the ice nucleating abilities of Kaolinite particles appeared to improve with decreasing temperature. The saturation at which ice was observed to form decreased at lower temperatures. Also, as the temperature decreased, the number of particles able to nucleate ice increased.  Chapter 5: Mineral Dust, 57 150 145 n  Water saturation Blank substrate Kaolinite  140 J 135-^ 130 ?125-1 120 J  II  115 110 105 100  II I 240  245  250  255  260  Temperature \ K  Figure 5.1  The RHj at which liquid water droplets or ice particles were first observed to form  on Kaolinite as the RHj inside the cell was slowly increased. The open symbols indicate that water droplets were first observed and the solid symbols indicate that only ice particles were observed with no indication of the formation of water droplets prior to ice formation.  Figure 5.2  Images of ice crystals and Kaolinite particles from two different experiments at  245.2 and 243.2 K . Cell temperature was decreased at a constant rate of 0.1 K min" until ice 1  crystals were observed to form.  Chapter 5: Mineral Dust, 58 5.2.2 Ice nucleation on Goethite The ice nucleating abilities of Goethite were determined by applying the  RH ramp t  experiment conditions. The RHj at which either water droplets or ice particles were observed are illustrated in Figure 5.3 for the blank hydrophobic glass slide as well as Goethite. The open symbols indicate that water droplets were first observed and the solid symbols indicate that only ice particles were observed with no indication of the formation of water droplets prior to ice formation. In the experiments involving Goethite particles at temperatures above ~ 252 K , water droplets were observed first (open symbols). Ice nucleation was not observed to occur at a later time.  The results in Figure 5.3 show that at temperatures above ~ 252 K , water droplets  condensed at water saturation, as expected. This observation leads to the conclusion that ice nucleation never occurred at temperatures above ~ 252 K and below water saturation for the current experimental conditions (observation time and Goethite particle concentrations). At temperatures below ~ 252 K ice nucleation was always observed. Images illustrating formation of ice on Goethite are shown in Figure 5.4. The formation of the liquid phase did not precede ice nucleation.  A s is shown in Figure 5.3, ice formed at RHj well below water  saturation. Comparable to Kaolinite, the ice nucleating abilities of Goethite particles appeared to improve with decreasing temperature.  At temperatures below 252 K ice was observed to form  close to ice saturation, and well below water saturation.  Chapter 5: Mineral Dust, 59 150 — o •  145 140  o  Water saturation Blank substrate Goethite  135 130 * 125  or 120 115 110 1054 100  240  245  250  255  260  Temperature / K  Figure 5.3  The R H ; at which liquid water droplets or ice particles were first observed on  Goethite as the RHj inside the cell was slowly increased. The open symbols indicate that water droplets were first observed and the solid symbols indicate that only ice particles were observed with no indication of the formation of water droplets prior to ice formation.  Figure 5.4  Images o f ice crystals and Goethite particles obtained from two different  experiments at 241.3 and 239.2 K . Here cell temperature was decreased at a constant rate of 0.1 K min" until ice crystals were observed to form. 1  Chapter 5: Mineral Dust, 60 5.3  Comparison with other results  5.3.1 Kaolinite Ice nucleating properties of Kaolinite were first investigated by  Roberts and Hallett  [1968]. Dust samples, consisting of up to ~ 10 particles, were placed on a microscope cold 4  stage under controlled RHj conditions. The authors noted that at temperatures of 254 K and above, it was always necessary to reach water saturation before the appearance of ice crystals. Below ~ 254 K deposition nucleation was observed on particles between 0.5 to 3 um in diameter. This observation is generally consistent with the results presented in this thesis. During RHj  ramp experiments involving Kaolinite, deposition nucleation was observed below ~ 252 K . In Figure 5.5, the current onset RHj for Kaolinite is compared with results obtained by Roberts  and  Hallett [1968]. Whereas Roberts and Hallett [1968] observed an approximately constant onset RHj at ~ 120% below 254 K , in the current data the onset RHj at which ice nucleates decreases with temperature. The discrepancy may be explained by the fact that Roberts  and Hallett [1968]  decreased the total number of Kaolinite particles used to nucleate ice at lower temperatures. Also, the Kaolinite sample employed by  Roberts and Hallett [1968] may have different  characteristics to that employed in the current study. During their investigation of ice habits, Bailey  and Hallett [2002] considered the onset of  ice formation on Kaolinite dust between 213 to 261 K . The corresponding data is plotted in Figure 5.5 and represents the nucleation onset at which less than 0.1% of the dust particles acted as IN. The data was obtained by adhering Kaolinite particles, 5 to 10 um in diameter, to a glass filament which was positioned inside a thermal diffusion chamber. The authors noted their data converged smoothly with earlier nucleation results of appears to be an offset of - 5 % RHj.  Bailey and Hallett [2002]. However, there  The reason for the difference in results obtained in the  current study and those presented by Bailey  and Hallett [1968] is unclear. However, the number  of particles used to coat the bare glass filament was not quantified by suggesting a different Kaolinite surface area may have been used.  Bailey and Hallett [1968],  Chapter 5: Mineral Dust, 61  140 Water saturation  1354  Our data N Roberts andtyllett [1968] 8a/fey and Ha/fetf [2002]  A o  130-4  125  4  s  1204  A  a:  MA  11 s;  1154  1104  1054  100230  235  240  245  250  255  260  Temperature / K  Figure 5.5  A comparison of current onset RHi results for Kaolinite with those of Roberts and  Hallett [1968] and Bailey and Hallett [2002]. Current data corresponds to the conditions at which water droplets (open squares) or ice crystals (solid squares) were observed. Roberts and Hallett [1968] observed immersion freezing above 254 K and deposition nucleation at lower temperature on particles between 0.5 to 3 pm in diameter.  The threshold of nucleation activity  was taken as the appearance o f one ice crystal in ~10 particles. The results from Bailey and 4  Hallett [2002] correspond to the onset of several ice crystals on Kaolinite particles (between 5 to 10 pm in diameter) adhered to a glass filament.  5.3.2  Chapter 5: Mineral Dust, 62 Comparison of current data with previous mineral dust studies other than Kaolinite Several different studies have shown that mineral dust can efficiently nucleate ice below  water saturation and in the deposition mode. Here, current Kaolinite and Goethite results are  [Mohler et al., 2005b; Archuleta et al., 2005; Knopf  compared with previous mineral dust studies  and Koop, submitted]. Mohler et al. [2005b] investigated ice nucleation on Arizona test dust (ATD) particles between 194 and 24IK. A T D is composed of a mixture of different minerals, but mainly consist of silicates, calcite, and clay minerals.  Dry particles, ~0.1 to 1.5 pm in  diameter, were injected directly into a large A I D A aerosol chamber. A s is shown in Figure 5.6, A T D particles were observed to nucleate ice at low supersaturations in the deposition mode. The ice nucleating abilities of commercially available aluminium oxide silicate (3Al203:2Si02), and iron oxide (Fe 03) were investigated by 2  (AI2O3),  alumina-  Archuleta et al. [2005].  The authors considered particles that were 50, 100, and 200 nm in diameter.  Only the ice  nucleation results for alumina-silicate and iron oxide particles are plotted in Figure 5.6.  Archuleta et al. [2005] noted that all particle types were as effective or better at initiating ice formation as compared with homogeneous freezing conditions.  Second, the ice nucleating  properties of mineral dust were observed to improve at lower temperatures and for larger particle sizes. In the study by  Knopf and Koop [submitted] the ice nucleating properties of A T D were  investigated between 197 to 293 K and below liquid water saturation. A T D particles, 0.7 to 10 pm in diameter, were deposited on a hydrophobic Herasil quartz plate positioned inside a chamber which operated in the Knudsen regime. Below 240 K ice nucleated on A T D particles via deposition nucleation at RHj well below water saturation. This onset RHi data for A T D is plotted in Figure 5.6. The results from the previous studies corresponding to the ice nucleating abilities of various types of mineral dust in the deposition mode are plotted in Figure 5.6.  The data  illustrates that the ability of mineral dust to nucleate ice is controlled by its composition, size, and morphology. Several conclusions can be drawn from these observations. First, all studies have found that, in general, mineral dust is clearly an active I N below water saturation in the deposition mode. Second, the ice nucleating properties of mineral dust are size dependent. This was conclusively shown by  Archuleta et al. [2005]. Finally, the variety of different types of  mineral dust in the atmosphere makes the task of quantifying the overall ice nucleating ability of  Chapter 5: Mineral Dust, 63 mineral dust difficult. Hence, more systematic studies are required to quantify the ice nucleating properties of these particles. A logical start would be to determine the  J  hel  and the contact angle  for each type of mineral dust component. This could set the basis for modelling or predicting I N of authentic atmospheric samples. In the subsequent section of this thesis the ice nucleating properties of Kaolinite and Goethite are quantified in this manner.  170 s  \ A  160J •  150  T  •  140J  i 130  120-1  110  100  210  215  220 Temperature / K  Water saturation Goethite, our data Kaolinite, our data Kaolinite, Roberts and Hallett [1968] Kaolinite, Bailey and Hallett [2002] Arizona Test Dust, Mohler et al. [2005b] Arizona Test Dust, Knopf and Koop [submitted] • 500nm, 100nm, 50 nm 3AI 0 :2Si0 , Archuleta era/. [2005] 2  A  Figure 5.6  3  2  500nm, 100nm, 50 nm Fe 0 , Archuleta era/. [2005] 2  3  A comparison of the current measurements of the onsets of ice nucleation of  Kaolinite and Goethite particles with data for other mineral dust measured in previous work.  Chapter 5: Mineral Dust, 64 Determination of onset J  5.4  for mineral dust  h e l  The above results are used to determine the heterogeneous nucleation rate coefficient, J  h e l  ,  for mineral dust in the deposition mode. In the deposition mode,  number of ice nucleation events, available surface area, and time.  J  h e t  depends on the  During the  RHi  ramp  experiments involving Kaolinite and Goethite either water droplets or ice crystals were observed to form first, depending on temperature.  Hence, two different approaches for calculating  J  h e l  were applied, depending on whether ice or water was observed first. In the experiments involving Kaolinite and Goethite at temperatures above ~ 252 K , water droplets condensed at liquid water supersaturation on four separate occasions. When ice did not nucleate first, the ice nucleating properties o f mineral dust were quantified by determining the upper limit to heterogeneous nucleation rate coefficient,  J  ,  v p h el  in an approach  similar to that involving soot. The upper limit to the heterogeneous nucleation coefficient for Kaolinite and Goethite was determined with Eq. 4.1 by taking r = 60 seconds, x = 0.95, and 4 ; = geometric surface area of the dust particles. The J  corresponding to the specific experimental  up el  conditions is summarized in Table 5.1 and Table 5.2. Below ~ 252 K ice nucleated first on Kaolinite and Goethite.  The heterogeneous  nucleation rate coefficient for experiments in which ice was observed to form first may be calculated in the following manner he,=-^r  [Salcedo et al., 2001; Salcedo et al., 2000],  ,  J  Eq.  5.1  where co is the number o f ice crystals nucleated, r is the observation time, and A is the total s  surface area available for heterogeneous nucleation. A t the onset o f ice nucleation co is equal to unity.  The observation time for this single nucleation event was 20±10 seconds.  The total  surface area, A , ranged from 3.3T0 to 2.5T0 nm . These values are listed i n Table 5.1 and 4  5  2  s  Table 5.2. To calculate A a geometric surface area for the clay particles was assumed. A n s  example of a single Kaolinite dust sample is shown in Figure 5.7. The figure also includes an average size distribution of eleven Kaolinite samples used in  RHj  ramp experiments.  Correspondingly, an example of a single Goethite dust sample is shown in Figure 5.8. The resulting values of  Jhet,  with estimated upper and lower limits, are listed in Table 5.1 and Table  5.2 for Kaolinite and Goethite, respectively. considering the error in A and r. s  Chapter 5: Mineral Dust, 65 The uncertainty in J was determined by het  Chapter 5: Mineral Dust, 66  (a)  90 in  C  o  b 80-1 1 5 70-i i fjj 60d)  (b) 1404:  0) o  ••ro e  30  " 20 o c  «  I? 10-1 i0  16  MUM 20  25  30  35  40  45  50  Diameter / microns  Figure 5.7  (a) A n image of the bottom of the cell showing the deposition of ~ 300 particles  of Kaolinite on a hydrophobic slide. The surface area of the particles is (6.5±0.8)T0 pm . 4  2  (b) Average size distribution of eleven different Kaolinite samples. Geometric mean diameter is 8.0 pm; geometric standard deviation is 1.2; surface area mean diameter is 19.7 pm (parameters defined by Reist [1993]).  Chapter 5: Mineral Dust, 67  Table 5.1  The heterogeneous nucleation rate coefficient for eleven separate Kaolinite  experiments.  Parameters describing experimental conditions include onset temperature and  onset RHj at which either water droplets or ice crystals were observed to form first. The geometric surface area o f each Kaolinite sample is also included. The uncertainty in Jhel was evaluated by considering error in the observation time (20±10 sec) and the surface area of particles. Here the lower and upper limits to Jhe, are J ™t and J et, respectively. The table also l  up  lists the total number of ice particles observed to form during the course of the experiment.  (cm"sec"')  Total number of ice particles formed  5±1  98"  0  122±4  6.5±0.8  77"  0  124±3  . 5±1  98"  0  250.2±0.2  114±5  7±2  38<71<175*  1  5  249.1±0.2  114±3  8±1  37 < 62 < 142 *  1  6  247.0±0.2  109±3  3.4±0.5  87 < 149 < 352 *  1  7  245.2±0.2  109±3  12±3  23 < 43 < 114*  1  8  243.2±0.2  103±3  4.4±0.1  69 < 113 < 255 *  4  9  240.7±0.3  103±3  16±3  18 < 32 < 73  10  240.4±0.2  103±3  24±4  12 < 20 < 48*  40  11  239.1±0.3  104±3  4±1  64 < 114 < 276 *  10  Temperature (K)  RHj (%)  1  254.8±0.2  120±3  2  252.6±0.2  3  252.5±0.2  4  Experiment  Surface area (•10 urn ) 4  2  rlow ° hei ^  j u  hel ^  jup het  J  2  6  25  In Experiments no. 1, 2, and 3 ice nucleation was not observed; water droplets nucleated first, hence J was determined by employing Eq. 4.1. The confidence level in j"h et is 95%. p  In Experiments no. 4 - 11 ice nucleated first, hence J , was determined by using Eq. 5.1. he  Chapter 5: Mineral Dust, 68  (a)  (b)  5  10  15  20  25  30  35  40  45  50  Diameter / microns  Figure 5.8  (a) A n image of the bottom of the cell showing the deposition of ~ 1000 particles  of Goethite on a hydrophobic slide. The surface area of the particles is ( 1 9 ± 4 ) T 0 pm . 4  2  (b) Average size distribution of six different samples of Goethite. Geometric mean diameter is 6.1 pm; geometric standard deviation is 1.2; surface mean diameter is 17.8 pm (parameters defined by Reist [1993]).  Chapter 5: Mineral Dust, 69  Table 5.2  The heterogeneous nucleation rate coefficient  for six separate Goethite  experiments.  Parameters describing experimental conditions include onset temperature and  onset RHj at which either water droplets or ice crystals were observed to form first. The geometric surface area of each Goethite sample is also included, noting that experiments no. 1,5, and 6 were performed on the same sample. The uncertainty in J  was estimated by considering  hel  the error in the observation time (20±10 sec) and the surface area of the particles. Here the lower and upper limits to Jhet are J ™, and J"h el, respectively. The table also includes the total number of 1  p  ice particles formed during the course of the experiment.  Experiment  Temperature (K)  RHj (%)  Surface area (•10 um ) 4  2  J  hel ^ het J  >  J  het  (cm^sec )  Total number of ice particles formed  1  1  256.9±0.3  119±5  19±4  0  2  250.1±0.3  118±5  10±2  27" 28 < 50 < 118"  3  247.0±0.4  108±6  18±3  15 < 27 < 66*  2  4  244.6±0.5  110±7  5.0±0.7  57 < 99 < 227 *  2  5  241.3±0.2  102±3  19±4  15 < 27 < 65*  6  6  239.2±0.3  102±2  19±4  15 < 27 < 65*  10  1  In Experiment no. 1 ice nucleation was not observed; water droplets nucleated first, hence J"  p h el  determined by employing Eq. 4.1.  In Experiments no. 2 - 6 ice nucleated first, hence J , was determined by using Eq. 5.1. he  was  Chapter 5: Mineral Dust, 70 5.5  Contact Angle In order to determine the impact of the measured ice nucleating properties of Kaolinite  and Goethite on the formation of the ice phase in lower tropospheric clouds, the ice nucleating behaviour of these mineral dusts must be characterized in terms of the contact angle. This step is necessary because it allows Jhet to be extrapolated to higher and lower saturations. The following discussion focuses on the calculation of the contact angle of an ice embryo formed on the surface of Kaolinite and Goethite between 239 and 258 K . Following the steps outlined in Chapter 4.4, experimentally-determined Jhet values (Chapter 5.4) have been used to determine AG* , acl  r , and 6  for each Kaolinite and Goethite experiment and are listed in Table 5.3. The uncertainty in the contact angle was determined by considering the uncertainty in the temperature (~ ±0.2 K , depending on the experiment), RHj (arising from uncertainty in the temperature and the frost point), surface area (depending on the experiment), interfacial tension (102 to 111 mJ m" ), and observation time (20±10 sec). 2  The contact angle pertaining to each of the experiments is plotted as a function of temperature in Figure 5.9. In the figure, solid symbols correspond to experiments in which ice particles nucleated first, with no indication of the liquid phase prior to ice formation; open symbols correspond to the limiting value of the contact angle for experiments in which water droplets were observed to form first. In the latter experiments in which water droplets were observed to form first at liquid water saturation, only the lower limit to the contact angle was determined. For the experiments in which ice was observed to nucleate first, a line of best fit is included. The linear regression between ~240 to 250 K for the contact angle of Kaolinite and Goethite is: ^Kaolinite  = -230.9+ 1.0 T  Eq. 5.2  Goethite = -276.7+ 1.2 T  where T is the temperature in Kelvin.  Eq. 5.3  The standard deviation, determined from the linear  regression, is 1.4° and 1.9° for Kaolinite and Goethite data, respectively.  Chapter 5: Mineral Dust, 71 Table 5.3  Onset parameters pertaining to the formation of a single spherical-cap ice crystal  in the Kaolinite and Goethite ramp RHi experiments for which Jhet was determined. AG* , was C  evaluated by using the pre-exponential term, A, equal to {10 }cm" sec"'. The radius of the 26  2  critical-size ice germ was evaluated by assuming no lattice strain in Eq. 2.20. Finally, the contact angle was determined by using Eq. 4.4. The uncertainty in J , AG *, r, and 6 was hel  act  evaluated by considering uncertainty in the temperature, onset RHj, surface area, interfacial tension, and observation time.  Onset Temp  Onset RHj  (K)  (%)  1  254.8±0.2  120±3  98"  1.94±0.03  11  22*  2  252.6±0.2  122±4  77°  1.94±0.03  10  23*  98"  1.93±0.03  9  24*  Experiment  J '  •Jhel  0 M  \G  V Jhet T ^ ^ Thet "P J  (cm' sec ) 2  _1  (10  r* -<r*<r* (nm) low  acl  1 9  J)  up  ef <e< ef" ow  (degree)  Kaolinite  3  252.5±0.2  124±3  4  250.2±0.2  114±5  38 < 71< 175  5  249.1±0.2  114±3  6  247.0±0.2  7  1.92±0.03  12 < 15 < 34  12 < 18 < 21  37 < 62 < 142  c  1.92±0.03  12< 15< 19  16 < 19 < 21  109±3  87<149<352  c  1.87±0.03  17 < 23 < 34  12 < 15 < 17  245.2±0.2  109±3  23 < 43 < 114  1.90±0.03  17 < 23 < 34  12 < 15 < 17  8  243.2±0.2  103±3  69 < 113 < 255  1.85±0.03  . 30 < 62 < 300  4 < 9 < 12  9  240.7±0.3  103±3  18<32<73  c  1.88±0.03  34 < 78 < 220  0 < 8 < 12  10  240.4±0.2  103±3  12<20<48  c  1.89±0.03  32 < 68 < 220  0<8 < 13  11  239.1±0.3  104±3  64<114<276  1.82±0.03  30 < 52 < 170  5<9< 13  1  256.9±0.3  119±5  27"  2.01±0.03  11  22*  2  250.1±0.3  118±5  28 < 50< 118  1.93±0.03  9 < 12 < 17  17<20<24  3  247.0±0.4  108±6  15<27<66  1.93±0.03  16 < 28 < 110  8 < 15 < 18  4  244.6±0.5  110±7  57 < 99 < 227  1.87±0.03  12<21<70  8 < 15 < 20  5  241.3±0.2  102±3  15 <27<65  1.88±0.03  39 < 92 < 220  0<7<12  6  239.2±0.3  102±2  15<27<65  1.87±0.03  28 < 88 < 220  0<8<14  c  c  c  c  Goethite  0  If water droplets were observedfirst,Jhe"  p  c  c  c  c  c  was calculated with Eq. 4.1.  * The lower limit to the contact angle was determined from Jhet • P  c  If ice nucleated first, Jhet was determined based on Eq. 5.1.  Chapter 5: Mineral Dust, 72  238  240 242 244 246 248 250 252 254. 256: ' 258:  Temperature / K Figure 5.9  Contact angle of an ice crystal formed under heterogeneous nucleation conditions  vs. temperature. Here a spherical-cap ice germ was assumed to form on a flat insoluble Kaolinite or Goethite surface (geometric mean diameter equal to 8 pm and 6.1 pm) respectively. Neglecting lattice strain, the contact angle was determined for N\ = 3.1 T O  28  m" and o-j = 106.5 3  v  mJ m" by using Eq. 4.3 and Eq. 4.4. Solid symbols correspond to experiments in which ice particles nucleated first, with no indication of the liquid phase prior to ice formation. Open symbols correspond to the limiting contact angles for experiments in which water droplets were observed to form first. For experiments in which ice was observed to nucleate first, the data is fitted to reflect the dependence of the contact angle on temperature. For Kaolinite and Goethite at temperatures between -240 to 250 K, respectively.  Q aoUmte K  = -230.9 + 1.0 T a n d  6 oethite G  = -276.7 + 1.2 T,  Chapter 5: Mineral Dust, 73 5.6  Atmospheric implications  In order to determine the impact of the measured ice nucleating properties of Kaolinite and Goethite on the formation of the ice-phase in lower tropospheric clouds, one must determine the maximum number of ice particles that can be produced under atmospheric conditions as a function of temperature and saturation. In order to do this Jhet as a function of temperature and saturation is required. However, the present experiment setup is only sensitive to a narrow range of  J  h e l  values. In each of the RHj ramp experiments, the temperature and saturation were  ramped until either the liquid phase or the ice phase was observed to form. As soon as the saturation and temperature were such that  J  was observed. For atmospheric purposes  h e t  J  was equivalent to this narrow range, nucleation h e l  must be extrapolated to higher and lower  supersaturations by using classical nucleation theory. Figure 5.9 illustrates that the contact angle, 6, decreases with temperature for both Kaolinite and Goethite. Between ~240 to 250 K, the relationship between the contact angle and temperature is described by Eq. 5.2 and Eq. 5.3 for dxaoiimte and dGoethue, respectively. Assuming 8 is independent of saturation, the contact angle is used in Eq. 2.22 to extrapolate  J  h c l  to higher  and lower saturations (below RH = 100%) for temperatures 240, 245, and 250 K. W  For example, at 240 K the contact angle for a nucleation of a single ice crystal on Kaolinite is 8.3° (Eq. 5.2). Assuming 8Kaolinite is constant with saturation,  J  h e l  at water saturation  (i.e. RHj = 138%) is equal to 7T0 cm" sec"' according to Eq. 2.23. This value is determined by 25  2  assuming o is 106.5 mJ m" and the density of ice is 0.92 g cm" . Predicted J 2  n  3  h e t  is plotted as a  function of saturation in Figure 5.10 (Kaolinite) and 5.11 (Goethite). The uncertainty in  J  h e t  was  determined by considering the standard deviation in dKaoiimie and dcoeMte from the fit to the data in Figure 5.9.  Chapter 5: Mineral Dust, 74 1E264  Figure 5.10  •  1  '  '  •  '  .„.!  :  '  '  3  The heterogeneous nucleation rate coefficient vs. supersaturation for Kaolinite.  Each curve corresponds to a different temperature: 240, 245, and 250 K. The contact angle at each temperature was determined by applying the parameterization Eq. 5.2. Assuming that the contact angle is independent of saturation, the nucleation rate coefficient (Eq 2.22) at each temperature is plotted as a function of supersaturation. Each J  het  100%. Dotted curves indicate the upper and lower bounds for J  hel  in OfCaolmite-  is extrapolated up to RH = W  as a result of the uncertainty  Chapter 5: Mineral Dust, 75 1E27j  .  i  .  !  .  1  •  1  •  1  «  1  >  1  1E26J  UH  1.00  '  1 -]  pJ—i  1.05  .  1  1.10  •  1  1.15  •  1  1.20  1  1  1  i  1  1.25  1.30  1.35  1.40  s Figure 5.11  The heterogeneous nucleation rate coefficient vs. supersaturation for Goethite.  Each curve corresponds to a different temperature: 240, 245, and 250 K . The contact angle at each temperature was determined by applying the parameterization Eq. 5.3. Assuming that the contact angle is independent of saturation, the nucleation rate coefficient (Eq 2.22) at each temperature is plotted as a function of supersaturation. Each 7 100%. Dotted curves indicate the upper and lower bounds for in Ocoethite-  J  t e  is extrapolated up to R H  h e l  as a result of the uncertainty  W  =  Chapter 5: Mineral Dust, 76 The number o f ice particles that can be produced under atmospheric conditions can be determined by modifying Eq. 4.5 to consider the number density of mineral dust instead of soot, such that  »«*="*,/[ - P(-- to4' )] 1  /  ex  T  E  q- 5  4  where nice is the number density of ice particles produced (litre" ), ndust is the number density of 1  1  2  mineral dust (litre" ) A is the surface area of a single dust particle (cm ), and r is the total time p  required to form a cloud (60 minutes). The number o f ice particles produced by Kaolinite and Goethite will vary with Jhet and bulk surface area of dust in the atmosphere. A large quantity of data has been published on mass density of mineral dust in the atmosphere [Husar et al, 1997;  Perry et al, 1997; Prospero, 1999; Falkovich et al, 2001; VanCuren and Cahill, 2002; Sassen et a/.,2003; Clarke et al, 2004] and it ranges from 0.2 to 2000 pg m" . It has been found that the 3  density of mineral dust varies with location, altitude, and season. Goudie and Middleton [1992] analyzed long-term meteorological records for a large number of regions and concluded that there is no single global pattern of dust-storm frequency trends. For this reason, the dust mass densities pertaining to specific topographical locations are not considered in this thesis.  Instead,  a mass density of ~ T10" g m" (close to the lower limit of mass density of mineral dust in the 6  3  atmosphere) is employed to estimate the number density of ice particles as a function of saturation and temperature. A n example calculation pertaining to Kaolinite is considered here. For ice nucleation in the atmosphere ndust and/lp are assumed to be 91.8 litre" and 1.3-10" cm . A value of 91.8 litre" 1  1  7  2  for ndust was calculated by assuming a dust radius of 1 pm, Kaolinite bulk density of 2.6 g cm" , 3  a geometric surface area, and a dust mass concentration (consisting of Kaolinite particles only) in the atmosphere of ~ 1T0" g m ' . A value of 1.3-10" c m for A was calculated based on a 6  3  7  2  p  geometric surface area of spherical dust particles of 1 pm mean radius - a typical radius of atmospheric dust particles [Ginoux et a/.,2001; Prospero 1999; Falkovich et o/.,2001; Blanco et  al, 2003; Perry et al, 1997; Petit et al, 2005; Jickells et al, 2005]. With these assumptions, n  ice  for temperatures between ~ 240 to 250 K could be determined in combination with the Jnet values plotted in Figure 5.10 and Figure 5.11. For example, at 240 K and an extrapolated relative humidity of 138% (equivalent to water saturation) « , i s 91.8 litre" for Kaolinite. In the case of 1  ce  Goethite, nice was determined by taking s Goethite bulk density of 4.27 g cm" . 3  Chapter 5: Mineral Dust, 77 The number density of ice produced as a function of saturation for Kaolinite and Goethite is plotted in Figures 5.12 and 5.13, respectively for temperatures of 240, 245, and 250 K. The uncertainty in n accounts for the standard deviation in the contact angle. In both cases, the ice  number density of ice particles increased with a decrease in temperature and increasing saturation. For each experiment «, is extrapolated up to R H = 100%. Both figures include the ce  W  Meyers et al. [1992] parameterization (Eq. 4.6) which predicts the total number of active ice nuclei in the atmosphere at a function of supersaturation, in the deposition mode and condensation freezing mode. If dust mass density is ~ 1T0" g m" , at lower temperatures both 6  3  Kaolinite and Goethite particles initiate the nucleation of enough ice crystals to account for the number of ice nuclei typically found in the atmosphere. This suggests that the contribution of mineral dust to formation of ice at conditions relevant to the lower troposphere is significant. It is interesting to note that the curve describing typical atmospheric number densities of ice nuclei [Meyers et a/., 1992] is significantly different from the trend predicted by classical nucleation theory. The current results consider the ice nucleating properties of Kaolinite and Goethite exclusively, whereas Meyers et al. [1992] considers the ice nucleating properties of atmospheric aerosols which are composed of a mixture of species. Each mineral type or other solid insoluble particle will have its own nucleating onset RHj. Hence the continuous curve of Meyers et al. [1992] may reflect this non-homogeneity of atmospheric particles.  Chapter 5: Mineral Dust, 78  100 90  80-  •240 K  CD  245 K 70-  250 K  Meyers etal. [1992]  Q.  8  1c  60-  5  i  0  •8 40H 9  -Q  E 304 20  A  10J 0. 1.00  Figure 5.12  1.10  1.05  ' i  1  1.15  1.20  1.25  1.30  1.35  1.40  s  Number density o f ice vs. saturation for Kaolinite assuming mass density of dust  in the atmosphere is 1 pg m" . The number density of ice is a function of the nucleation rate 3  coefficient determined in Figure 5.10. The solid curves correspond to different temperatures o f 240, 245, and 250 K for which the contact angle was 8.3°, 13.3°, and 18.3° respectively. Corresponding dotted curves reflect the standard deviation of 0Kaoimite. The  Meyers et al. [1992]  parameterization predicts the total number of active I N in the atmosphere at a function of supersaturation, in the deposition mode and condensation freezing mode.  Chapter 5: Mineral Dust, 79 60  50  8.  J  404  CD O  O  £ . 30 CD  ~o Cl)  204  3 C  10-1  240 K 245 K 250 K  Meyers et al. 1.00  Figure 5.13  [1992] 1.40  Number density of ice vs. saturation for Goethite assuming mass density of dust  in the atmosphere is 1 pg m" . The number density of ice is a function of the nucleation rate 3  coefficient determined in Figure 5.11. The curves correspond to different temperatures of 240, 245, and 250 K for which the contact angle was 7.6°, 13.5°, and 19.4° respectively. Corresponding dotted curves reflect the standard deviation of  6 emeGo  The  Meyers et al. [1992]  parameterization predicts the total number of active IN in the atmosphere at a function of supersaturation, in the deposition mode and condensation freezing mode.  Chapter 5: Mineral Dust, 80 5.7  Conclusions  The results presented in this section establish that Kaolinite and Goethite nucleate ice between 239 and 258 K , below water saturation, and in the deposition mode. In this study mineral dust particles were exposed to conditions resembling those of the lower troposphere. A t temperatures above ~ 252 K , water droplets were observed to form first. This shows that at these temperatures ice nucleation does not occur below water saturation under the current experimental conditions.  However, below ~ 252 K ice particles were consistently observed to form well  below liquid water saturation. This suggests deposition nucleation on Kaolinite and Goethite particles occurs in this temperature range. Further, the ice nucleating abilities of Kaolinite and Goethite improved with decreasing temperature. A s the temperature decreased, ice nucleated at relative humidities well below water saturation. Combining all the previous results and the current study, it appears that below water saturation at temperatures below ~ 252 K , ice nucleation on many types of dust particles is efficient  [Bailey and Hallett, 2002; Roberts and Hallett, 1968; Archuleta et al, 2005; Mohler et  al, 2005b; Knopf and Koop, submitted]. In order to quantify the ice nucleating abilities o f Kaolinite and Goethite, the heterogeneous nucleation rate for each of the experiments was determined.  Also, classical  nucleation theory was employed to determine the contact angle of an ice embryo formed under  RHj ramp experiments. The contact angle was determined by  conditions defined by the specific  assuming a single spherical-cap ice germ formed from supercooled water vapour on a flat and insoluble Kaolinite or Goethite surface.  In both cases, the contact angle was observed to  decrease with temperature. In the current study, at lower temperatures ice particles were observed to form well below water saturation. For this reason the determined  J  hel  quantified the ice nucleating ability of  Kaolinite and Goethite at the onset of nucleation defined by specific temperature and RHj. In order to determine the atmospheric implications of this data,  J  hel  was extrapolated to higher  relative humidities by assuming that the contact angle does not vary with saturation. allowed for the extrapolation of and 250 K .  J  hel  This  as a function of saturation and temperature between ~ 240  This relationship was used to evaluate the implication o f dust as I N in the lower  Chapter 5: Mineral Dust, 81 troposphere under typical atmospheric conditions. For mineral dust densities of 1 pg m" the 3  values were often significantly higher than those predicted by  n  ice  Meyers et al. [1992], suggesting  this type of mineral dust may significantly alter the formation of the ice-phase in lower tropospheric clouds at temperatures below ~ 252 K , and below water saturation.  Chapter 6: Summary and conclusions, 82  Chapter 6  Summary and conclusions  6.1  Ice nucleating properties of soot and dust  In the previous chapters, the ice nucleating properties of several types of soot and mineral dust were investigated between 239 to 258 K , below water saturation, and in the deposition mode. For comparison, this data is plotted in Figure 6.1. In  RHj ramp experiments soot or dust  particles were deposited on the bottom surface of the flow cell; RHj inside the cell was increased, and the conditions under which water droplets or ice crystals formed were determined with a reflected-light microscope. In the case of soot (N-hexane soot, Lamp Black 101, Degussa FW2, and Printex 40) water droplets always nucleated first at temperatures above 248 K .  This  observation lead to the conclusion that ice nucleation does not occur below water saturation and in the deposition mode at these experimental conditions. Below ~ 243 K ice particles were occasionally observed first. However, since ice formed close to liquid water saturation, ice may have formed through condensation freezing. This thesis also considered whether the ice nucleating abilities of soot would change following exposure to atmospherically-relevant quantities of ozone.  However, even after an  exposure equivalent to 80 ppb O3 at atmospheric pressure for 13.7 days, the ice nucleating abilities of Lamp Black 101 were not different from results obtained for the unexposed samples.  RHj ramp experiments were performed on Kaolinite and Goethite dust between 239 and 258 K .  Below ~ 252 K , ice particles consistently formed first, demonstrating deposition  nucleation is the dominant mode of ice formation for these experimental conditions. This data also showed that the ice nucleating ability of Kaolinite and Goethite improves with a decrease in temperature.  This conclusion was drawn from the observation that, as the temperature  decreased, ice formed at lower supersaturations. The above observations may be explained by considering the composition and morphology of soot and mineral dust particles. Soot has an onion-like structure which consists of a graphite-like core coated with an amorphous organic layer (Figure 1.4). This amorphous coating may lack the ordered structure necessary to be a good I N .  Chapter 6: Summary and conclusions, 83 150  1454 1404  1354  1304 125  X 120  115  1104 1054  100 240  245  250  255  260  Temperature / K Figure 6.1  A comparison of the R H i at which water droplets or ice crystals were observed to  form first for  RH ramp experiments involving soot, Kaolinite, and Goethite. Solid symbols t  correspond to experiments in which ice crystals were observed to nucleate first; open symbols correspond to experiments in which water droplets condensed first.  In comparison, the nanostructure of Kaolinite and Goethite is crystalline. Kaolinite is built of alternate layers of Si20s and Ab(OH)4 in such a way that hydroxyl groups are exposed on one surface and silica on the other  [Mason, 1960; Buol et al, 2003]; Goethite consists of  Fe(III) that is surrounded by three O and three O H groups to give FeOOH an octahedral structure. Ice nucleating abilities of Kaolinite may be credited to the presence of the hydroxyl groups which allow water molecules to bind to the surface. Since the arrangement of O H groups  Chapter 6: Summary and conclusions, 84 on basal plane of Kaolinite complements the hexagonal ice lattice, the ice lattice will have coherent geometry (Chapter 2.4) and form on the substrate surface by means of H-bonding. In a similar manner, the arrangement of the OH groups in Goethite also favours ice nucleation.  6.2  The heterogeneous nucleation rate coefficient  The ice nucleating abilities o f soot and mineral dust were quantified by determining the heterogeneous nucleation coefficient at specific experimental conditions.  experiments involving soot the average J"  p h el  For  RHi  ramp  was determined to be 50 cm" sec"'. This value was 2  calculated by assuming ice nucleation did not occur during an observation time of 60 seconds, at a confidence level of 95%, and over a known surface area of soot. Since soot expressed poor ice nucleating abilities between 243 and 258 K , the constant RHj a better constraint on  J"  p  in the deposition mode.  h el  experiment was invoked to provide  During this experiment n-hexane soot  (air/fuel = 2.4) was held at 247.5 K and below liquid water saturation. Since water or ice did not nucleate during the 8 hour experiment, determined J"  p h t  was 0.1 cm^sec' , much smaller than the 1  previous result of 50 cm" sec" . 2  1  In order to quantify the ice nucleating abilities of Kaolinite and Goethite, the heterogeneous nucleation rate for each of the experiments was also determined. If water droplets were observed first, J  up  ranged from 77 to 98 cm" sec" at temperatures above ~ 252 K , at liquid 2  et  1  water saturation. For experiments in which ice nucleated first  J  het  was calculated by assuming a  single nucleation event took place over 20 seconds at a specific RHj and temperature. For  J  Kaolinite, Goethite  6.3  J  values ranged from 20 to 149 cm" sec"' between 103 to 114% RHj, whereas for 2  hel  ranged from 27 to 99 cm" sec"' between 102 to 118 % RHj, respectively. 2  hel  Contact angle comparisons  In order to further quantify the ice nucleating abilities o f soot and mineral dust, heterogeneous nucleation theory was employed to determine the contact angle of the ice embryo  Chapter 6: Summary and conclusions, 85 formed on the insoluble particle. The contact angle obtained for soot, Kaolinite, and Goethite is compared in Figure 6.2. The lower limit to the contact angle was determined for several RHj  ramp experiments involving soot particles. Since water droplets or ice crystals were always observed to nucleate at liquid water saturation, the contact angle reflects this dependence on saturation. As the temperature decreased, water droplets or ice crystals were observed at higher supersaturations.  Hence the lower limit to the contact angle increased at lower temperatures  because it was constrained by liquid water saturation. The contact angle for the deposition nucleation of an ice germ on Kaolinite and Goethite was also determined. The ice nucleating properties of Kaolinite and Goethite improved below ~ 252 K (i.e. ice was observed to form at lower relative humidities). Consequently, the contact angle decreased with temperature and saturation. equations:  6  Kaohme  = -230.9 + LOT and 0  Goelhile  This relationship was fitted with two  = - 2 7 6 .7 + 1.2T for temperatures between ~  240 to 250 K .  6.4  Atmospheric implications  For soot, the upper limit of 0.1cm" sec" was used to determine the number density of ice 2  formed at 124% RHi.  1  B y applying elemental carbon mass densities representative of specific  scenarios, it was determined that the number density of ice particles formed as a result of J^  t  was significantly less than the number of IN observed in field studies  [Meyers et al., 1992]. This  lead to the conclusion that soot alone cannot account for the presence of the ice-phase in lower tropospheric clouds, below water saturation, and above ~ 243 K . Kaolinite and Goethite exhibited strong ice nucleating properties at temperatures below ~ 252 K . Since the current experimental setup was only sensitive to a narrow range of Jhel values, heterogeneous nucleation theory was used to predict  J  hel  at higher or lower supersaturations by  assuming the contact angle is independent of saturation.  The relationship between  J  het  and  saturation was used to determine the number density of ice particles formed at conditions typical to the lower troposphere. If the mineral dust mass density is 1 pg m" , the number density of ice 3  particles formed as a result of Jhel is significantly larger than the number of I N observed in field studies  [Meyers et al, 1992]. This lead to the following conclusion: Kaolinite and Goethite may  Chapter 6: Summary and conclusions, 86 play an important role in formation of the ice-phase in lower tropospheric clouds, below water saturation, and below ~ 252 K .  30 X  2826-  X  24-  X  22-  D  </>  CD  0  xo x  20X X  CD  o  18 J  0)  16  n 144  Kaolinite o  6 44  A  2  X  Goethite best fit Soot - constant RH. experiment' Soot - ramp RH. experiments 4  238 240 242 244 246 248 250 252 254 256 258 260 Temperature / K Figure 6.2  Contact angle as a function of temperature for  Kaolinite, Goethite, and  constant  RHi  RHj  ramp experiments with soot,  experiment with soot.  Solid symbols correspond to  experiments in which ice crystals were observed to nucleate first; open symbols correspond to experiments in which water droplets condensed first.  References, 87 References Akhter, M . S . , A . R . Chughtai, and D . M . 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