Basaltic glasses from Iceland and the deep sea: Natural analogues to borosilicate nuclear waste-form glass.
MicliMlJ.J«rcinovfc and Rodney C.Ewing
D«c«mb«r,1987 BASALTIC OLACSBI FROM ICBLAHD AVD THE DB» SBA: ITOBAX. AMALOGUBf TO BOROflLICATB MUCL1AB WMT1-F0RM GLASS
Michael J. J«rcinovic and Rodn«y C. Ewing
D«c«mb«r, 1987
D«parta«nt of Geology Th« University of New Mexico Albuquerque, New Mexico USA 87131 11
list of Tables iv list of Figures vi Suenery xiii Abstract xvi 1 introduction 1 1.1 Alteration 6 1.1.1 Palagonitizaticn 6 1.1.2 Palagcnitizaticn Rates 9 1.1.3 Secondary Mineralization 13 1.2 Samples 21 1.2.1 Iceland 21 1.2.2 Dredge Sanples 26 1.2.3 Drill Core Samples 26
2 Techniques 29 2.1 Thin Section Preparation 29 2.2 Scanning Electron Microscopy 31 2.3 X-Ray Diffraction 31 2.4 Electron Microprobe Analysis 32 2.5 Analytical Electron Microscopy 34
3 Results 35 3.1 Icelani 35 3.1.1 General cements 35 3.1.2 Fresh Mater Alteration 36 3.1.2.1 Pleistocene Snhjiftrtnl Volcanic» 37 3.1.2.1.1 Palagonite 37 3.1.2.1.2 Cssentation 42 3.1.2.2 Tungufell 55 3.1.2.2.1 Palagcnite 55 3.1.2.2.2 Oawntation 68 3.1.3 Seawater Alteration 72 3.1.3.1 General Conomts 72 3.1.3.1.1 Palagcnite 76 3.1.3.1.2 Cementation , 92 11.' 3.2 EKedge Sau&m 107 3.2.1 ROagonite 107 3.2.2 OsasntiiHin 117 3.3 Erill Om Saaples 128 3.3.1 ffelagcnite 128 3.3.2 t 3.4 Analytical Electron Microscopy 144 3.4.1 Saaple Description 144 3.4.2 Analytical Ilectrcn Micxceoopy 147 3.4.2.1 OSMI 113521-69 147 3.4.2.2 UGM1 113715 153 3.4.3 conclusion* 156
4 Discussion 158 4.1 ROagonite 158 4.2 Secondary Mineral Authigenasis, Solution Concentrations, and Mass Balance 180 4.3 Alteration Rates 200
5 Conclusions 206 5.1 Corrosion Machanisn 206 5.2 Alteration Products 207 5.3 Mass Balance 209 5.4 Alteration Rates 210 Acknowledgements 212 213 iv
Table 1. Senary of palagcritizaticr. rates. Table 2. Secondary minerals associated with altered basaltic glasses. Table 3. Samples fzen Iceland. Table 4. Deep sc\ dredge sanples. Table 5. Deep e-» ir'H oare samples. Table 6. Apr Tvtrit pnlagonite rind thicknesses in Icelandic samples. Table 7. Average microprobe analyses of glass and corresponding palagonite rinds frar Icelandic subglacial vcicanics. Table 8. Average njryoprobe artalyses of clays in Icelandic subglacial volcanics Table 9. Average oicroprobe analyses of zeolites from Iceland. Table 10. Average riToprobe analyses of glass and corresponding palagonite rinds from hyaloclastitas and pillov basalt tins at Tungufell, Iceland. Table 11. Average ndcroprcbe analyses of clays from Tungufell, Iceland. Table 12. Average mlcroprobe analyses of glass and corresponding palagonite rinds frca seawater altered nyaloclastites and pillow basalt rims in Iceland. Table 13. Average Bicroprob* analyses of clays in Icelandic seawater altered nyaloclaatif.es and pillov basalts. Table 14. Apparent palagonite rind '.'hicknesses in deep sea samples. Table 15. Average v tcroprobi anaJLyse» of glass and corresponding palagonite rinds in dv«p sea dredge tuples. Table 16. Results for constant Fa in the high-Fe, low-si deep sea dredge pal»gonite«. v Table 17. Average miexoprobe analyses of days from deep sea dredge and drill
Table 18. Average aiexoprobe analyses of deep sea zeolites. Table 19. Average microprobe analyses of glass and corresponding palagonite rinds in deep ssa drill core sasples. Table 20. Compositions of phases A, B, and C of OSMf 113521-69 determined by MM. Table 21. Gompositions of different areas of U9tf 113715 determined by AIM. Table 22. Relationship between pH and aluninum spaciaticn. Table 23. ftolagonitization rates inferred for Iceland samples from apparent r«M thickness and corresponding sample age. Table 24. Palagonitization rates inferred for dredge samples from apparent rind thickness and corresponding sample age. Table 25. Palagonitization rates inferred for drill core samples frcro apparent rind thickness and corresponding sample age. VI LIST Of 7IGDRBS
Figure 1. Schematic representation of a subglacially-produced table mountain (after Jones, 1970; Allen, 1980b). Figure 2. Location nap of Iceland. Figure 3. Summary of characterization scheme. Figure 4. Riotonicrograph of a palagonitized glass frcn Icelandic subglacial volcanic ridge (Trolladyngia). Figure 5. Glass normalized-palagcnite composition of Icelandic subglacial volcanics. Figure 6. Variation plots of microprobe analyses of M-l clays. Figure 7. Variation plots of microprobe analyses of M-3 clays. Figure 8. Variation plots of microprobe analyses of M-6 clays. Figure 9. Variation plots of microprobe analyses of M-7 clays. Figure 10. Variation plots of microprobe analyses of Mid.-l clays. Figure 11. Variation plots of microprobe analyses of Mid.-3 clays. Figure 12. Variation plots of microprobe analyses of R-l clays. Figure 13. Variation plots of microprobe analyses of R-2 clays. Figure 14. Plot of exchangeable cations in zeolites in sample M-6 (Mosfell, Iceland). Figure 15. Sketch of a thin section of Tungufell, Iceland sample T-2. Figure 16. Glass nonnalized-palagonite composition of samples T-3 and T-4 from Tungufell, Iceland. Figure 17. Results of a microprobe traverse of a palagonite rind in sample T- 2, Tungufell, Iceland. This is Traverse 1 of 10 which were done on rinds associated with zeolite filled vesicles. vii Figure 18. Results of a microprobe traverse of a palagonite rind in sample T- 2, Tungufell, Iceland. This is Traverse 2 of 10 which were done on rinds associated with zeolite filled vesicles. Figure 19. Results of a microprobe traverse of a palagonite rind in sample T- 2, Tungufell, Iceland. This is Traverse 8 of 10 which were done on rinds associated with zeolite filled vesicles. Figure 20. Results of a aicroprobe traverse of a palagonite rind in sample T- 2, Tungufell, Iceland. This is Traverse 9 of 10 which were done on rinds associated with zeolite filled vesicles. Figure 21. Results of a microprobe traverse of a palagcnite rind in sample T- 2, Tungufell, Iceland. This is Traverse 2 of 6 which were done on rinds associated with vesicles filled with unoriented clay. Figure 22. Results of a microprobe traverse of a palagonite rind in sanple T- 2, Tungufell, Iceland. This is Traverse 4 of 6 which were done on rinds associated with vesicles filled with unoriented clay. Figure 23. Comparison of microprobe traverses of the two palagonite rind types in sample T-2, Tungufell, Iceland. Figure 24. Comparison of microprobe analyses of T-2 days, Tungufell, Iceland. Figure 25. Variation plots of aicroprobe analyses of T-3 clays. Figure 26. Variation plots of microprobe analyses of T-4 clays. Figure 27. Potassium content of zeolites (from microprobe analysis) vs. depth from the pillow surface in sample T-2, Tungufell, Iceland. Figure 28. Relationship between apparent palagonite rind thickness and vesicle size in a pillow rin from seawater altered sample Reyk.-l, Reykjanes, Iceland. viii Figure 29. Glass nornalized-palagonite ocopositions of seawater altered sanple Ko.-l, Kopavogur, Iceland. Figure 30. Glass rcrmalized-palagonite ccqposition of seawater altered samples fran Iceland. Figure 31. Results of a microprobe traverse of a palagonite rind in sanple Br.-2, Brianes, Iceland. This is Traverse 1 of 6. Figure 32. Results of a microprobe traverse of a palagonite rind in sanple Br.-2, Brumes, Iceland. This is Traverse 2 of 6. Figure 33. Results of a ndcroprobe traverse of a palagonite rind in sanple Kb.-l, Kopavogur, Iceland. This is Traverse 2 of 6. Figure 34. Results of a microprobe traverse of a palagonite >*ind in sanple Ko.-l, Kopavogur, Iceland. This is Traverse 4 of 6. Figure 35. Results of a microprobe traverse of a palagonite rind in sample Kb.-l, Kopavogur, Iceland. This is Traverse 5 of 6. Figure 36. Results of a microprobe traverse of a palagonite rind in sanple Ko.-l, Kopavogur, Iceland. This is Traverse 6 of 6. Figure 37. Results of a microprobe traverse of a palagonite rind in sample A- 1, Arnarnesvogur, Iceland. This is Traverse 1 of 6. Figure 38. Results of a microprobe traverse of a palagonite rind in sample A- 1, Arnarnesvogur, Iceland. This is Traverse 4 of 6. Figure 39. Results of a microprobe traverse of a palagonite rind in sanple Reyk.-l, Reykjanes, Iceland. This is Traverse 2 of 6. Figure 40. Results of a microprobe traverse of a palagonite rind in sample Reyk.-l, Reykjanes, Iceland. This is Traverse 4 of 6. Figure 41. Clay and zeolite conpositior» (microprobe analyses) in sample Ko.- 1, Kopavogur, Iceland. ix Figure 42. Variation plots of ndcroprobe analyses of Br.-l clays (Brimnes, Iceland). Figure 43. Variation plots of nicroprobe analyses of Br.-2 clays (Brisnes, Iceland). Figure 44. Variation plots of microprobe analyses of A-l clays (Arnarnesvogur, Iceland). Figure 45. Variation plots of microprobe analyses of Reyk.-l clays (Reykjanes, Iceland). Figure 46. Variation plots of microprobe analyses of Reyk.-2 clays (Reykjanes, Iceland). Figure 47. Scanning electron photomicrographs of zeolites in sample v~.-l, Ropavogur, Iceland. Figure 48. Scanning electron photomicrograph of a) chabazite and b) calcium- silicate in sample A-l, Arnarnesvogur, Iceland. Figure 49. Plot of exchangeable cations in zeolites in sample A-l (Arnarnesvogur, Iceland). Figure 50. Scanning electron photomicrograph of phillipsite in sample Br.-2, Brismes, Iceland. Figure 51. Plot of exchangeable cations in zeolites in sanple Br.-2 (Brimnes, Iceland). Figure 52. Glass nornalized-palagonite ccnpositions of low FeO, high SiO2 deep sea dredge samples. Figure 53. Glass nornalized-palagonite compositions of high FeO, low SiO2 deep sea dredge samples. Figure 54. Results of a microprobe traverse of a palagonite rind in deep sea dredge sample USNM 113715 (low FeO, high SiC^). This is Traverse l of 6. X Figure 55. Results of a ndcroprobe traverse of a palagcnite rind in deep sea dredge saqple USNM 113715 (low PsO, high SiOj). This is Traverse 2 of 6. Figure 56. Results of a nicroprobe traverse of a palagonite rind in deep sea dredge sanple US»! 113715 (low FeO, high SiOg). This is Traverse 4 of 6. Figure 57. Results of a microprobe traverse of a palagonite rind in deep sea dredge sample USNM 113521-69 (high Fed, low SiOj). This is Traverse 3 of 6. Figure 58. Results of a microprobe traverse of a palagonite rind in deep sea dredge saqple USNM 113521-69 (high FeO, low SiO2) • This is Traverse 4 of 6. Figure 59. Results of a micxoprobe traverse of a palagonite rind in deep sea dredge sample USNM 113521-69 (high FeO, low SiC^). This is Traverse 5 of 6. Figure 60. Si<>2 variations across palagonite rinds in sanples USNM 113715 (Traverse 4) and USNM 113521-69 (Traverse 4). Figure 61. Plot of exchangeable cations in phillipsite in deep sea dredge sanple USNM 113487-3. Figure 62. Glass nonnalized-palagonite ccnposition of deep sea drill core saqples. Figure 63. Results of a microprobe traverse of a palagonito rind in deep sea drill core sanple ODP 37 332B 6-2. This is Traverse 1 of 6. Figure 64. Results of a microprobe traverse of a palagonite rind in deep sea drill core sample ODP 37 334 17-1. This is Traverse 1 of 6. Figure 65. Results of a microprobe traverse of a palagcnite rind in deep sea drill core sasple ODP 37 334 17-1. This is Traverse 6 of 6. Figure 66. Results of a microprabe traverse of a palagonite rind in deep sea drill core savple
During the five phases of the JSS project, a substantial data base has been generated to describe the corrosion of nuclear waste borosilicate glass. The experimental data have been interpreted by a model, GLASSOL, developed by Grambow (JSS Report 87-02), and the sane model has been used to assess the long-term performance of the OOGEMA glass, JSS-A, in an idealized granitic repository for various flow rates and glass surface areas (Grambow et al., in press). To validate the long-term extrapolation of the performance of the COGEMA glass using GIASSOL, a detailed study of natural, basaltic glasses — from Iceland and the sea floor — was completed. Previous workers (Lutze et al. 1985; Malow, lutze, and Euring 1984) have demonstrated the analogous corrosion behavior of basaltic and borosilicate glasses under experimental conditions. The main purpose in studying basaltic glasses was to: (l) identify long-term corrosion products; (2) to determine the corrosion rate and the effect of silica-concentration on the corrosion rate; and (3) to determine major element mass balance as a function of the type of alteration system.
This report provides a detailed analysis of the alteration process and products for natural basaltic glasses. Information of specific applicability to the JSS project include:
* The identification of typical alteration products (Table 2) which should be expected during the long-term corrosion process of low-silica glasses. Of particular importance are the results of the detailed examination of "amorphous" leached layers using analytical electron microscpy (AEM). xiv These AIM results (Figures 12 through 76) demonstrate that the leached layers are not amorphous, but contain a relatively high proportion of crystalline phases, mostly in the form of smectite-type clays and that channels through the layer provide immediate access of solutions to the fresh glass/alteration layer interface. Thus, glasses are not "protected" from further corrosion by the surface layer.
* Although the corrosion rate is not easily extracted from natural samples because the contact time of the solution with the glass is difficult to establish exactly, there is compelling evidence that corrosion proceeds with two rates — an initial rate, k+, in silica-undersaturated environments and a long-term rate, k, in silica-saturated environments (Grambow et al. 1986). This is important because it not only supports the fundamental premise of the GIASSOL model, but also demonstrates that there is no unexpected change in corrosion rate over long periods of time. The long-term corrosion rate (an estimated minimum value of 0.01 microns/1000 yr, see Figui.% 92) of natural basalt glasses exposed to sea water on the open ocean floor (5°C) is consistent with that of borosilicate glasses. Note, corrosion rates must be corrected for differences in the chemical composition of the glasses and temperature differences in the alteration environments of natural basalt glasses as compared to experimental conditions for the borosilicate glasses (see Grambow et al. in press). XV * Detailed examination of reaction products in vesicles has provided examples of the effect of precipitating phases on silica concentration in the solution (closed systan conditions). As predicted by the GIASSOL model, precipitation of silica-containing phases can result in increased alteration of the glass as manifested by greater alteration layer thicknesses (Figure 15). This emphasizes the importance of being able to predict which phases form during tfct reaction sequence.
* For natural basaltic glasses the flow rate of water and surface area of exposed glass are critical parameters in minimizing glass alteration over long periods of time» Mass loss of principal elements is highest under high flow rate conditions, and alteration layers are, in many cases, prominent on even the narrowest fractures. This is most dramatically demonstrated by different samples of basalt glass which are of the same age and altered (5°C) on the ocean floor. In some cases, fresh, unaltered glass remains, and in others, the glass is completely altered. The long-term stability of basalt glasses is enhanced when silica concentrations in solution are increased.
In summary, there is considerable agreement between corrosion phenomena observed for borosilicate glasses in the laboratory and those observed for natural basalt glasses of great age. xvi ABSTRACT Basaltic glasses from subglacial-volcanic deposits in Iceland, and from submarine deposits along the Icelandic coast and the sea floor were analyzed to determine alteration effects. The primary alteration product is palagonite which forme rinds on all exposed glass surfaces, regardless of environment or age. Palagcnite replaces the glass from the cuter surfaces toward the interiors of hyaloclastite grains or pillow rims, forming a reddish, vitreous, hydrolysis product. Petrographic evidence suggests that palagonit i nation is isovolumetric, with the possible exception of some high-Fe, low-Si deep sea dredge palagonites. In subglacial volcanic palagonites and most seawater formed palagonites, the rind is ccnposeci of colloidal material resen*0 ing ncntronite. Although palagonite appears to be amorphous by Scanning Electron Microscopy (SIM) and optical microscopy, Analytical Electron Microscopy (AEM) reveals considerable microstructure. AEM results indicate zoning in terms of the degree of crystallinity, microstructures, and chemical composition within palagonite rinds. Crystal dimensions do not exceed 50 nm. At least three phases have been identified in palagonite rinds by AEM. Palagcnite rinds vary considerably in composition from one location to another, with Al content being an important environmental indicator. Seawater-formed palagonite bears the chemical signature of seawater with respect to alkali content (Na2O from .2 to 2.7 wt. t and K20 from .14 to 3.62 wt. % in seawater-formed palagonite compared to <.2 wt. % and <.l wt. % respectively for fresh water formed palagonites). High K-content suggests that deep sea palagonites may be more illite-like. Deep sea dredge samples are unique in that many are highly enriched in Fe (enriched by a factor of as much as 5 above the amount in the glass progenetor) in the form of Fe- oxyhydroxide. This Fe-enrichment may result from alteration near sea-floor xvii hydrothermal systems. Ti-depletions oacur in sane palagonites relative to parent glass, most ocnmcnly in deep sea dredge samples. The cause of Ti depletion has not been determined. There is no major ocnpositional difference between deep sea dredge (low Fe) and drill core samples, tlais palagonitization of drill oore glasses nay have occurred in the open ocean environment prior to burial.
Compositional gradients across palagonite rinds are more strongly developed in low Al, zeolite-associated palagonites. Al and Ca tend to Increase toward the interior (later-formed) portions of the rinds. Fe and Mg decrease correspondingly. These gradients are interpreted as being the result of changing solution conditions (e.g., increasing pH) as hydrolysis of the glass proceeds under closed system conditions. Seme compositional zoning is also present in sane rinds not associated with zeolites. This zoning is similar in direction but lower in magnitude compared to zoning in rinds associated with zeolites.
Depletion of Al in the palagonite rind relative to parent glass is necessary for zeolite formation. The zeolite formation results fran high reaction progress, an indication of high pH. Zeolites form in nearly closed systems when flow rates are low. The presence of zeolites correlates with thicker palagonite rinds. Relatively thick rinds result fran high dissolution rates at high pH or long solution exposure tines. Zeolites may not form even if significant Al is released during palagonitization if the ratio of the effective solution volume to the volume of glass reacted is high. Saponite precipitation proceeds zeolite if the rate of release of Mg+Fe exceeds Ca release during palagonitization. Pnillipsite is not caiman in zeolitized xviii
Iceland sawoles but is present in JCM? &*ep sea samples. Chabazite is the daninant zeolite in all laolitized samples except those from the deep sea. Analcime, when present, is the last zeo?.its to form, due to the consumption of K and Ca by earlier zeolites and closed system conditions. Mass balance calculations indicate that little mass loss occurs when zeolites are formed and that the elements required to form autMgenic phases in zeolitized samples may originate entirely from glass dissolution. Sn fresh water, alkalis and alkaline earths may show some net system loss ever after clay and zeolite formation. Clay cement is often the result of high mass input under open system conditions. Net mass loss of Si, Mg, Ca, Na and K are high if clays or zeolites do not form.
Palagonite rind thickness does not correlate with age due to differences in solution contact time as compared to the true age of the sample. In addition, reaction rates vary depending on the solution pH and silica concentrations in solution. Minimum rates are between .01 to 1 microns/1000 yr. There are not any consistent differences in rind thickness between dredge and drill core samples, as one might expect due to differences in silica concentrations in interstitial waters as compared to those in the open ocean. TICM
This report is the contribution of the University of New Mexico to Fhase V of the JSS project. The report assesses the feasability of using natural analogues (e.g., basaltic glasses) in the prediction of the long-term alteration behavior of borosilicate nuclear waste form glasses and to describe alteration products and processes in natural basaltic glasses from selected geochemical environments.
Short term glass corrosion processes can be investigated via laboratory lear* tests. Geochemical models (e.g., GIASSOL) must be used to extrapolate such test results to repository-relevant time periods (103 to 105 years). Naturally occurring glasses have been exposed to aqueous environments for similar time periods (and up to perhaps as much as 107 years). Such glasses can be useful analogues if corrosion processes in natural glasses and nuclear wasteform glasses are similar. This similarity in alteration behavior has been demonstrated for natural basaltic glass and nuclear waste form borosilicate glass by laboratory experiments (Malow, Iutze, and Ewing 1984; Lutze et al. 1985; Byers, Jercinovic, and Ewing 1986) and extended to a comparison of surface layers (palagonites) formed as alteration products on basaltic glasses (Iutze et al. 1985).
Information from natural analogues regarding alteration phase types, amounts, and compositions are used to provide input into, and evaluate results of, geochemical modelling. Modelling requires the choice of the appropriate alteration phases. Certain phases might be excluded as possibilities if they are not identified in natural systems. If corrosion models can be effectively 2 applied to the natural systems, that is, if modelling can predict observed alteration phenomena, then natural glasses should provide a reasonable basis for model verification (Ewing and Jercinovic 1987).
In order to use natural systems as analogues, the alteration history of the selected deposits, including constraints on starting conditions must be known. This alteration history is obtained by field observations of deposits and microanalyticzJ. characterization of samples. Types and compositions of alteration products are used to evaluate the geochemical conditions at the time of formation, i.e., the alteration environment. The alteration environment can be described in terms of starting solution composition (pH, seawater, fresh water, etc.), temperature, pressure, and whether the system is open or closed (e.g., flow rate). The system mass balance can be assessed if phase compositions and amounts are known. An open system might result in net mass input while little mass loss would be expected if the system were closed. Natural geochemical systems are rarely truly closed.
The information obtainable from naturally altered glasses can be placed into four major categories:
CORROSION MECHANISM - Before the effects of alteration can be compared from one glass to another or from one environment to another, the actual process by which alteration takes place must be understood. Observations of glasses altered in experiments and in nature can help to constrain the possible processes. Questions pertinent to the alteration mechanism include: 1) is hydration a dominant mechanism in the aging of basaltic glasses when exposed to water (as it has been demonstrated in low silica glasses) or do hydrolytic 3 reactions daninate; 2) if dissolution occurs, is it congruent or inccngruent; 3) what is the role of diffusion; 4) is ion exchange an important step in the overall alteration process; and 5) how is the process affected by silica saturation? Natural samples can also be used to evaluate whether or not the processes change over long periods of time (thousands or millions of years).
ALTERATION HCDUCIS - Most of the data gathered relevant to glass corrosion are on the alteration products. There are five types of pertinent information: 1) phase identity; 2) phase composition; 3) relative proportions of phases; 4) alteration paragenesis; and 5) thermodynamic stability vs. the kinetics of formation of phases. Additionally, other properties of certain phases are important (e.g., density or ion exchange capacity). Once baseline data are established, the similarities and differences between samples from different environments or ages can be assessed. The effects of cementation processes on the hydrologic characteristics (e.g., flow rate) of the deposits can also be evaluated, of particular importance to thermodynamic reaction path modelling is the characterization of "amorphous", metastable reaction products which contain ultra-fine grained multi-phase precipitates of nonuniform crystallinity. A critical question involving dissolution/reprecipitation reactions in basaltic glassss is whether there is any volume change in the process. This is crucial in determining the amount of glass involved in the reaction, and along with density, is required for mass balance determination. Rdagonite density is poorly known and is certainly oompceitionally dependent. Accurate methods for determining palagonite density must be developed. 4 MASS BALANCE - The amounts and compositions of alteration products can be used to evaluate mass balance in natural alteration systems. Elements released during glass dissolution are either used in reprecipitation reactions or remain in solution, to be eventually removed from the system. Mass balance analysis will show which elements are lost (if any), the magnitude of such losses, and the circumstances under which element losses take place. Comparing many samples from a number of alteration environments and representing different ages and glass compositions may reveal the circumstances required for mass loss or the circumstances where material input has taken place. Sane environments may consistently offer the most closed system or open systan conditions. For instance, alteration taking place in the unsaturated (vadose) zone of a deposit will undergo temporary saturation and high flow rates during periods of high rainfall. Material released into solution will be rapidly carried away and the system nay be considered qpc-r.. In contrast, alteration in a stagnant lake may offer more nearly closed system conditions as flow rates should be low.
CORROSION RAIES - Rates of reaction may be estimated by observing many samples representing a wide range of ages. However, simply relating alteration rind thickness to age is not a reasonable procedure for defining absolute corrosion rates in naturally altered samples, as the actual contact time with the solution may not be known. Also, the age determinations for geologic deposits are often not well constrained. The contact time of glass with solution is a function of intermittent flow (i.e., in unsaturated zones), change in environment due to tectonic, climatic, or other influences, or a change in hydrologic characteristics of the deposits themselves (e.g., sealing due to authigenesis). The determination of reaction rates from rind thicknesses is 5 further complicated by the fact that the rates of reaction may have changed with changing conditions, the result of variations in pH or silica activity. Such factors may be evaluated in natural samples if the geologic setting and history of deposits are well known.
In this study we will compare alteration products and processes in basaltic glass samples from two selected areas: Iceland and the sea floor, ttie glasses occur in the form of glassy-fragmental deposits (hyalodastites) or pillow lavas. Pillow lavas are characterized by the pillow structure of the deposits, consisting of partially flattened, meter-sized, globular masses of basalt. The outermost centimeter of individual pillows is compos»** almost entirely of glass while the interior of the pillows is crystalline. If the effective E^QQ at the vent is equivalent to a water depth of approximately 200 m or more, pillow basalts are formed (Allen, Jercinovic and Allen 1982). At lower values of PH2O the combination of vesiculation from devolitilization and quenching can shatter the magma, resulting in hyalcclastites. Hyaloclastites may also form by spallation of pillow rims (Fisher and Schmincke 1984), a process that may be responsible for formation of most deep sea hyalcclastites.
Samples represent various alteration environments and ages. Most samples from Iceland were altered in fresh water with the exception of some collected near the stioreline. One of the major aspects of this study is a comparison of seawater altered samples from recently emerged shoreline areas of Iceland, dredge samples from the sea floor surface, and drill core samples from the oceanic crust beneath sediment cover. 1.1 AI3ERATI0N
1.1.1 PAIAG0NITIZA1TCN
Palagonite was a tenn first used by W. Sartorius von Waltershausen (1845) to describe the brown material in the groundmass of a tuff front Palagonia, Sicily. It has since taken on a variety of meanings ranging from reference to entire deposits (Pjetursson 1900) to microscopic materials associated with sideromelane glass (Peacock 1926; Peacock and Fuller 1928). More recently, the term palagonite has principally been used with regard to the generally vitreous or gel-like, optically isotropic, x-ray amorphous, reddish to yellow- brown pseudomorphic alteration product of sideromelane glass (e.g., Nayudu 1962; Hay and Iijima 1968a, Allen et al. 1981).
The physical and chemical properties of palagonite have been studied by a number of authors (e.g., Kay and Iijima 1968a 1968b; Honnorez 1972; Jakobsson 1972; Singer 1974; Fumes 1978 1980 1984; Fumes and El-Anbaawy 1980; Allen et al. 1981; Staudigel and Hart 1983). Palagonite has a refractive index which falls in the range 1.46 to 1.70 (Hay and Iijima 1968a 1968b; Stokes 1971) and a specific gravity in the range 1.93 to 2.42 (Hay and Iijima 1968a, 1968b; Fumes 1978). The glass/palagonite contact is sharp (Hay and Iijima 1968b; Geptner 1978) and is usually characterized by an etched or pitted glass surface (Hay and Jones 1972; Allen et al. 1981). In some cases, the glass exhibits penetration by microcnannels (Staudigel and Hart 1983) which may be equivalent to the "rootlet intergrowth" observed by Geptner (1978). 7 The mechanism by which palagonite forms has been proposed as a hydration process (Peacock 1926; Margenstein and Riley 1974), solid-state diffusion - hydration (Moore 1966), and solution precipitation (Hay and Iijima 1968b; Hay and Jones 1972). The solution - precipitation mechanism (also termed hydrogen -ion metasomatism [Hay and Iijima 1968b]) is based on the physical characteristics of altered samples (sharp glass/palagonite contact and an etched glass surface) as well as the chemical changes which result from palagonitization.
Ideas concerning the »i™*» and temperature of palagonitization range fran immediately with eruption at high (magnetic) temperatures (Bcnatti 1965) to time scales on the order of thousands of years after eruption at low (sedimentary) temperatures (Hay and Iijima 1968a 1968b). Jakobsson (1978) and Jakobsson and Moore (1986) have shown that palagonitization has occurred on Surtsey from temperatures as low as 40°C up to 150°C within several years after eruption. Higher temperatures increased the palagonitization rate there, a conclusion supported by the experimental work of Moore (1966) and Moore, Fornari, and Clague (1985). Fumes (1975) concludes from experimental palagonitization studies that:
". . . the process can be initiated as a post-eruptional (or post- cooling) phenomenon at any time when sideromelane encounters an aqueous environment., and that the initiation of the process does not require high temperatures."
Several general chemical characteristics of palagonite are evident from data published to date (e.g., Bonatti 1965; Hay and Iijima 1968a 1968b; Hay 8 and Jones 1972; Muffler et al. 1969; Hormorez 1972; Jakobsson 1972; Melson 1973; Singer 1974; Fumes 1978 1980 1984; Juteau et al. 1979; Furnes and El- Anbaawy 1980; Ailin-Fyzik and Scraner 1981; Allen et al. 1981; Staudigel and Hart 1983; Jakobsson and Moore 1986). The general conclusions that can be drawn from these studies include:
1) palagonitization is a ncn-isochemical alteration process and not siitple hydration;
2) the composition of palagonite is ccmnonly different from one occurrence to another;
3) chemically different palagonites can originate from chemically and physically similar glasses, emphasizing the importance of the alteration conditions (i.e., solution composition, temperature, flew rate).
Most major elements are depleted, to varying degrees, during palagonitization with the exception of iron, titanium, and aluminum (although even these elements may also be lost in seme specimens). Palagonites formed from high Fe and Ti glasses have Fe and Ti concentrations considerably higher than those palagonites formed from glasses lower in Fe and Ti (Furnes and El- Anbaawy 1980). Concentrations of immobile elements in palagonite rinds are thus greatly affected by their concentration in the parent glass. The behavior of the alkaline earth elements is quite variable and cannot be related to environment or glass composition. In all cases, the ratio Fe3+/Fe2+ increases from glass to palagonite. Alteration of basaltic glass in 9 the marine environment (e.g., Andrews 1977; Baragar et al. 1977; Bonatti 1965; Melson 1973; Pritchard et al 1978; Juteau et al. 1979; Staudigel and Hart 1983; Scarfe and Smith 1977) tends to produce palagonite which is enriched in potassium (palagonites containing 3 to 4 wt.% K20 have been produced from glasses containing less than 1 wt.% K20). Sodium can be retained to some degree during submarine palagonitization (glasses containing several wt.% Na2O alter to palagonites which have N^O values ranging from less than .5 to greater than 2 wt.%).
Summarizing the previous studies, the chemical composition of palagonite, is a result **:
1) the initial physical and chemical properties of the parent glass;
2) the properties of the aqueous solutions responsible for alteration; and perhaps,
3) the duration of alteration processes, which in rock- dominated systems can have an important effect on the properties of the aqueous solutions.
1.1.2 PAIAGONirrZOTGN RATES
In order to determine palagonitization rates, Moore (1966), Moore, Fornari, and Clague (1985), and Hekinian and Hoffert (1975) have compared 10 palagonite rinds formed in the deep sea (on basaltic glass dredge samples) with manganese coatings on the sane samples. Well constrained rates of Mn precipitation allowed the derivation of age/thickness relationships. Hekinian and Hoffer t confirmed their Mn thickness ages using fission track dating techniques on samples from the Mid Atlantic Ridge. Previous studies of palagonitization rates are summarized in Table 1.
Although Hekinian and Hoffert (1975) estimate palagonitization rates between 2.6 and 4.3 microns/1000 yr and Moore (1966) estimates rates to be in the range of 3 to 30 microns/1000 yr (at 5 to 50°C), Crovisier et al. (1986) indicate that both sets of data are actually not inconsistent and thus chose 2.6 to 4.3 microns/1000 yr as a reasonable rate. This agrees with the "hydration rind" formation rates of 3 microns/1000 yr found by Morgenstein and Riley (1974). Experimental dissolution of basaltic glass by Crcvisier et al. (1986) lead them to conclude that the dissolution rate is actually on the order of ISO microns/1000 yr at 3°C and 350 bars. Based on this and comparing their results to the palagonite growth rate of Hekinian and Hoffert (1975), they suggest palagonitization is a non-isovolumetric process.
Fumes (1975) studied palagonitization rates experimentally as a function of temperature and found that the rate increases rapidly with temperature and is non-linear with time. In experiments over 14 months, it was found that the rate of palagonitization of alkali basalt glass» is about .30 microns/month for the first 10 months from 20 to 70°C. Ihe palagonitization rate accelerates after 10 months at temperatures above 70°C, to about 3.3 microns/month. Jakobsson and Moore (1982 1986) have estimated palagonitization rates as a function of temperature based on detailed drill core studies of the Surtsey 11
Tabta 1. tiaaary of p»l»a,onl tint ton rata».
Condition» kafaranca
Natural palaaonitltatlon
3 aicrana/1000 yr taaoatar Norfanttain tnö II lay (1974)
2.é ta 4.3 aleran»/1000 yr •aavatar J°C •tkinian and »offert (1975)
3 ta 30 aferena/1000 yr tamatar 5 ta 50°C •aera (196A) Maara tt al. (19SS)
15 aUrana/1000 yr taanatar 48C Jtkabaion and Moor» (1966)
3 ta 20 aierera/1000 yr( fraih iiattr ind taaxatar Graabotf at »I. (198J)
.1 ailcrana/1000 yr (2) fraah ««ttr and taaitatar 6ra»*e« »t al. (198S) fipaHaantal palafenitflatien
.3 •icreni/Bonth fr»»h unr »nd Hintir Furna» (1975) 20 to 70°C
3.3 •ferant/aanth fr»»h Mtar »nd »aavatar Furna» (1975) •oov» 70sC »ftar 10 awnth*
100 aicrent/IOOO yr •aanatar Croviaiar »t al. (198S)
180 aricrana/IOOO yr •aawatir 3°C and 3S0 bar* Cravtaiar at al. C°86)
12 ta 17 •leront/Booth »act teluttan »t 200*C ••lov »t al. (1984)
rarvard rata. 12 tephra. Bie predicted rate from Surtsey is 15 microns/1000 yr for 4°C considering that the rate doubles with a rise in temperature of 12°C. Moore, Fornari, and Clague (1985) found the rate to increase 1.8X for a temperature rise of 10°C in seawater near Hawaii. Crovisier et al. (1986) have experimentally noted an increase in dissolution rate with pressure, obtaining values for 60°C of about 10 microns/yr at 1 bar and 14 nicrons/yr at 350 bars for tholeiitic glass in seawater.
Assuming a rate of 2 microns/1000 yr, centimeter-thick glassy rims on pillow basalt would not survive more than about five million years; and hyaloclastites (of average grain size less than 2nn), one »i "> lion years in the deep sea. However, fresh glass occurs in core samples of much greater age, such as DSDP Hole 335 [13 m.y. (Aumento et al. 1977)], 408 [20 m.y. (Iiiyendyk et al. 1978)], 407 [35 to 36 m.y. (Iuyendyk et al. 1978)], 382 [70 to 80 m.y. (Tucholke et al. 1979)], and 417 and 418 [105 to 110 m.y. (Donnelly et al. 1979) ]. Very old fresh water altered deposits, with much fresh glass remaining, also occur, such as the Miocene Frenchman Springs Flow of the Wanapum Basalts (12 to 14 million years old) in the Columbia River Basalt Group (Jercinovic, Ewing, and Byers 1986). Ancient deposits of submarine- produced hyaloclastites, now subaerially exposed, occur in Upper Triassic (200 to 210 million years) Hound Island Volcanics (Muffler et al. 1969) and the Carboniferous (about 340 million years) Iahn Dill deposits (Schmincke and Pritchard 1981). Clearly, assuming a single linear alteration rate is inappropriate. The effects of sealing of deposits by authigenesis in order to preserve ancient glasses has been proposed by a number of authors. Schmincke et al. (1978) suggest that sealing has preserved glasses in DSDP leg 46 deposits for the last 13 million years. Similarly, Muffler et al. (1969) 13 proposed that calcite effectively sealed the Hound Island volcanics from further alteration, allowing partial preservation of the glass. Sealing nay have also been an important factor in the preservation of glass in the Siletz River volcanics, which are 35 to 55 million years old (Keith and Staples 1985).
Grambow (1985) models the glass dissolution process as occuring with two rates: 1) an initial rate (silica undersaturated); and 2) a final rate (silica saturated). This has been substantiated (Grambow et al. 1985) by the examination of some naturally altered samples from British Columbia arid the deep sea. The forward rate is 3 to 20 microns/1000 yr and the final rate is about 0.1 microns/1000 yr. The long-term survival of glass samples can be explained by the approach to silica saturated conditions in the aqueous phase of well compacted samples (e.g., low flow rates) or those buried by sediments at the ocean floor.
1.1.3 SEOOMTARY MINERALIZATION
The rate and progress of palagonitization may be dependent on the precipitation of phases which decrease porosity and the access of water to the glass (Rimes 1974). Leached components must, therefore, be removed by percolating waters or held in solution for extensive palagonitization to occur (Fumes 1974 1975). The occurrence of secondary phases in altered glass deposits may be related to the conditions of alteration. The sequence of precipitation (paragenesis) provides a record of the chemical evolution of the system. 14
Berner (1980) defines an authigenic mineral as:
". . . one that forms within a sediment after burial, in other words, during diagenesis. It nay crystallize in the original pore space of the sediment or it nay fill pore space created by the dissolution of a pre-existing mineral. In the first case the s known as cenentrvticn and in the latter as replacement."
Falagonite is replacement-authigenic as it replaces the glass progenitor. Secondary minerals, such as zeolites and calcite, fill pore spaces and are, thus, oanentation-authigenic. Cementing materials affect the hydrologic properties (e.g. permeability) of the deposit in which they occur.
Table 2 lists the secondary minerals reported with palagonitized basaltic glass. Zeolites, clay minerals (particularly smectites), and calcite are, by far, the most cumiun secondary minerals associated with palagonites.
Among the zeolite minerals, analcime, phillipsite, and chabazite are reported most often. There is some evidence that zeolite type may be influenced by initial glass composition. Kristmannsdottir (1978) reports mordenite, heulandite, stilbite, and epistilbite in tholeiite host rocks, and chabazite, thomsonite, and mesolite-scolecite in olivine-tholeiite host rocks (lower bulk SiC^). Si/Al ratios are generally higher in phillipsites in marine sediments compared to those from mafic igneous rocks (Fisher and Schmincke 1984). Terrestrial phillipsite is generally more calcic (and less alkaline) than marine phillipsite (Fisher and Schmincke 1984). 15
T»blt 2. lacondary •trwrait »MOdtttd nith »(tared bataltfc ilttiti.
limril Oceurrtfict •ttpaetfv* (ofartnco*
Clay Mineral*
Saactltaa •antran ttfwilt* (.SCi.aa) 0Urad««(NT); stBrUl 31; 3,12,24,26 rctaatntto Cl; K; 1; OtOradfa; KNT) 1»; 7; 7; 21,1;19,20 OfBrtdft 21 unapaclflad taactfta OtOriIt; I(NT); CIS; 0«(J«) 16,27,28; 25; 2; 18 (andltaa Uallntta 42; CIS 5; 6 Illtta trout 111 it* *2; CIS 5; 6 Caiaawnita BtOr«dtt Chiartta Craup Chlarttt DtOrlll; KNT); SSOrt«|* 7; 33.29,23,19; 1.12 •raftnita C«JAIJ(OH)J»(JO10 KNT) 33,23,20 »popxyllttt OtOrill; OK») 12; 18 l»y«r CD 1 or(c«/a*nta>r1II oni t» KNT); GN 33; 30 KNT) 33 l*e*n 2ael •nalclaa Braup Analctaa IC; N; KNT); Cl; CIS; 7: 4,1,9,13; 33; 23; 2 OtOrUl; OKGan); M(II); 12,16; 17; 17; 18; t(NT); OfDradaa 14; 11 ualraifta KNT) 33,25 »•traitt« Grau» Natralita N; M Table 2. Secondary atnerala aesociatod with altered basat t ic gtaaeet (continued). Minaret Foraula Occurrence letpectlve References Chabaiite Croup Chaaeilte IC; I; CIS; DMritl; 7- 7; 2; 12,16; N; t(NT); 0t(«); S(NT) 8,9 ,13; 20; 18; U CHllnitt OWrlll 12 trfonlte N 4 •Mlllpalte Croup IC; OtOritl; H; CI; 7; 7,27,12.2t,J6: 4,8.9,13; 23 0«; CIS; AZ; CiS(Gon); I(HT) 23; 2,6; 4; 17; U ln* N 13 0S<«»n) 17 it t ta KMT); 0»(«) 2», 25,20; 18 Kaulendtte Croup Noulandlte CIS; KKT); OKS*) 6; 20; 18 KNT); Old»] Stllblte C«4 -281120 33. 20; 18 ClinaptUollte CIS; OS(Con) o; 17 Mordent te Group Morden!to IC«D; CIS; OI(SI) 33,29,20; »,• 18 Iplitllblte KNT) 20 'eujaeite Croup Faujaaite (» H 13 Unapoelflod 2eolite KUT); f. 25; 23 Cirbonatet Cateiia CaCOi Cl; IC; BSOrUt; H; T5.23; 7; 7,28,12,24,26,:6 4,8,9,10; OSOrodfO(MT); HKT); IC; CIS;31; 29,25,20; 23; 2,6; NIS; 0KSI); KMT); OSOrodfO 22; 18; 14; -.1 Ar«|ontte C«COj OSOriU; N; OSOradfa 28; 9; 11 •a»neaHe "fCO, OSDrill 28 SldarMa OSOrUI 26 Doloaite OSOrfll; SC; CIS 28; 23; 6 Antar t ta CIS 6 Caie-tlUcatai CyroUte OSOrUI; KMT) 12; 20 •ayarfta itnry 20 ZaopnvlUte Cas Ttbte 2. Secondary afnerel* tttoclated »itu altered beteltlc *lae*e* (continued). alnorel remit Occurrence •••pectiv* ••ftrmcti Hn-hydroxide OSOrill 21 ro*tiydroxlde OfOrill 26 fe«Mn hydroxide DSOrodae 11 Ffonttft M; OSOredte; OSOrilt 23; 11; 1* Un-Olid» M; OlOrtdit 2S; 11 My! Occurr«ne*« • ef vrinen « Atertt 1 Ian» 1*72 21 Molten and Thoacxon 1973 OC IrltitM CeliaMta 2 troy and Icnalnte 19U 22 Muffler et at. 19*9 Clt Cwwry It (and* 1 Col* and thoit 1963 23 Nevudu 19*4 Cl Celuabt* «lv«r ••••it 4 Coaan and I»in* 1»»* 24 »rltcherd, Cann and wood 1978 DSOrtdf* D*«p ••* dr«fl(« 5 rum** 19(0 25 •eblnton et tl. 1962 0Urtd|*(NT) Ot«p ••• drMf* (hydretntratl) 6 »urne* end 11 »r*a»»y 1980 2* Scarf* and laftft 1977 OfOr ill Ottp ••• drill e»r« 7 Gr****» *t at. 19U 27 Scftatncke 1963 0l(6*n) Ot«p »•• (fdfwrtl) a Kay and tljla* )»•*« 28 Scttafncke et al. K77 EN Colin Ntifhti (ItrMl) V nay and tljlae 19Mb 29 Sltvtldaaaon 19*2 H DIM i i 10 Nay and Jone» 1972 30 Sinfer 1974 NIS Wound i«l»nd volcanic* (Alttkt) 11 Neklnlen and Noffert 197S 31 Stake* and O'Noll 1962 1 Icilir4 (loa t«apor*turt> 12 NOAnorei 1971 32 Suaa*r* 197* KMT) Iceland (hydrother»»I) 11 lljlaa and atrad* 19*9 33 Toaatton end Krittaenntdsttir 1972 O« Oraton ceaatd Mtalta U jakob***n and Moore 19M ' OKU) Orefon (Illtlt River volcanic*) 15 JerclnovU, I«fn« *nd Iyer» 19M tc Santa Cruz Mountain* (CatHernia) 16 Juteau et al. 1979 » Glassy deposits affected by hydrothennal systans exhibit zeolite varieties which are restricted to specific temperature ranges. Mordenite, heulandite, stilbite, epistiibite, chabazite, thomsonite, and mesolite- scapolite are all found in deposits below 100°C and are replaced by laumontite at 100 to 120°C (Kristnannsdcttir 1978). Tcnasson and Kristuannsdottir (1972) report the complete disappearance of zeolites above 230°C. Nontronite, Fe-saponite, monGnorillonite, and chlorite are the most conmonly reported clays. Fe-saponite is the most prevalent clay in the submarine environment (Cole and Shaw 1983). The formation of nontronite and celadonite in the deep sea has been suggested to occur at low temperatures (0 to 20°C) and under oxidizing conditions (Rjrnusov et al. 1982). Isotopic evidence suggests that saponite forms from 30 to 200°C under "non-oxidative" conditions (Stakes and O'Neil 1982). The occurrence of chlorite is evidently restricted to glasses in the deep sea or to glasses that have been hydrothermally altered. Tomassan and Kristanannsdottir (1972) indicate that chlorite appears in hydrothermally altered deposits in Iceland between 230 and 280°C, with interlayered cUorite/montsnorillonite from 200 to about 230°C. This is consistent with observations in other hydrothermal alteration systems in Iceland (Kristmannsdottir 1975 1978). Calc-silicates are restricted to hydrothermally altered rocks but are only found at temperatures below 200°C (Kristmannsdottir 1978). Epidote, in contrast appears at temperatures above 200°C (Kristmannsdottir 1978). 20 Alteration paragenetic sequences (the relative tines at which minerals are deposited) have been defined fcr a number of palagonitized glasses. In Canary Islands deposits, Brey and Schmincke (1980) indicate the sequence (from the first mineral crystallized to last): chabazite-phillipsite- (analcime+clay [minerals listed in parentheses were precipitated at approximately the same time]). In deep sea drill core samples the sequence: sapaiite-gyrolite-phillipsite-chahazite-ginelin^ (opal+calcite) has been reported (Hannorez 1978). The melilite-nephelinite tuffs of Oahu, Hawaii exhibit these reported sequences: 1) phillipsite- chabazite-thansorute-gonnardite-ratrolite-analciffle- (laDntmorillonite+cpal+calcite) (Hay and Iijima 1968a); 2) phillipsite- chabazite-ar^cijiie-(iioTtinorillonite+<'-pal+calcite) (Hay and Iijima 1968b); and 3) K zeolite-Ca+Na zeolite-Na zeolite (Iijima and Harada 1969). One of the more detailed studies of authigenic mineral paragenesis in glassy basaltic rocks is that of the Siletz River volcanics of Oregon (Keith and Staples 1985). The following sequence is indicated: smectite-pyrite-calcite- thaisorute-ratrolite-ar^ciine-sajlecite-inesoli^^ apophyllite-^abazite-morderdte-^^dcite-laumontite-quartz. The alteration of the Siletz River volcanics is interpreted as occurring either relatively rapidly in seawater at 60 to 70°C (near a hydrothernal vent), as suggested by isotopic evidence, or in fresh water (meteoric water) at approximately 1O°C for a long period of time. With the exception of the Siletz River volcanics, phillipsite (K-rich zeolite) generally proceeds chabazite (Ca-rich zeolite) and analcime and calcite are late-stage cements. The occurrence of analcime and calcite after 21 calcic zeolites would be expected if PQ^ increased with time (Zen 1961; Hay 1966; Velde 1977). 1.2 1.2.1 ICELAND Iceland provides ubiquitous glass of different ages (years to millions of years) that have been altered in a variety of different conditions, such as seawater, glacial melt water, high temperature, or low temperature. Subglacial volcanism during the Pleistocene (Recent to 2 million years) produced large quantities of glass as table mountains (Figure 1), hyalodastite ridges (frcn fissure eruptions), and massive hyaloclastite deposits (Allen 1980a; Allen et al. 1982). Glass occurs as rims on pillow basalts and as hyaloclastite. locations of Pleistocene (< .7 m.y.) subglacial eruptives sampled include Trolladyngia, Hosfell, Middalsfjall, Raudafell, Blafell, and Valafell (Figure 2). Additionally, Plio-Pleistocene (.7 - 3.1 m.y.) deposits of basalt and hyaloclastite at Tungufell are also of subglacial origin (Saenundsson 1979). Seawater altered samples from Iceland were gathered at shoreline deposits near Reykjavik, including Brimnes, Kopavogur, and Amarnesvcgur, and near the southwestern-most point on Reykjanes Peninsula near Reykjanes (Figure 2). In all, twenty samples were included frcn Iceland, summarized in Table 3 (see also Byers, Jercinovic, and Ewing 1986). 22 topset, subaerial flows hyoloclastite tuff hyalodastite • basalt talus föreset, flow-foot bn-ccia Mgurt 1. Sehtitiitic representation et • lubjlacteUyprodueed t»blt mountain (after Jones, 1970; Alien, 1960b). 23 Figure 2. location mtp of Ictland. T»bl» 3. Staplat froa IcaUnd. location loaptt Invlronaant Itwrki TrollotfyntU Tr.-1 Uppar »lalttacana1 >ub«ltc KOtftll »•1 • tubtlacfal latal pil lo» unit N-3 tubflacial Hyaloclaatlta In ttrataevd K-6 • tubflactal Nyaloctaatlta tuff • •T • tubtlteial Myalaclatitta braeefa Irtamt lr.-1 •tafttocana, lata Karfna, »I»» »•U.-1 Uppar »laiitoeana3 Subflacfal Nyaloclattita tgfloaarata Mfd.-2 lue«(tcltl Croat bäddad conilaaartta Id.J luetlactal Myaloelaatita taaloaarata (audtftll 1-1 Subfttcial vitrie ath • -2 Subfltclat ittat pUlo» unit Utfall • t. -1 Pltfttoctn*3 fub«(aeial Nyatoclaatita tuff II.-4 • tub«laclat Ka froa) pil lo» frafawnt in tfaloaaratt Tunaufall T-2 ,7 to 3.1 »y3 Sub«l»eial7 »Ulo»t and inttrpillo» hyalaelattita T-3 m Subflaeltl? T-4 m Supflaclal? Ullo» r\m in flott Valtfall V-1 uppar »latttocana3 tue«(tc (opavofur ICO.-1 Malttocana, lata Karfn» a Arnarnatvofur »•1 »laiitaeana, ttta Marfna, aiiad Hyaloctattltt ttatoaarttt Intaralacfal* tayi.-1 Halttocan», lata Nar ^taaundtton tnd Ilnariton (1*90). ^Tryifvaten ind Jontion (1958). 'jtdannation, Jtkobtten, and Staaundtton (1977). 25 Subglacially erupted materials were initially exposed to glacial melt water (englacial lake) for an undetermined period of time. They have since been exposed to meteoric water, primarily through vadose zone percolation at relatively low temperatures (mean annual temperatures near 0°C). Localized high temperature (several hundred degrees) hydrothermal activity has also affected some deposits of subglacial volcanics. Samples from these areas were not included in this study. Samples from Briranes, Kopavogur, Arnarnesvogur, and Reykjanes are hyaloclastites and pillow basalts which are exposed along the Icelandic coast. These deposits evidently erupted beneath seawater when sea level was considerably higher than present (late interglacial or early post-glacial). Precise ages are unknown but most likely between 9000 years, when present day sea level was established (Saemundsson 1979), to 320,000 years, the beginning of the Gunz-Mindel interglacial (von Frechen and Lippolt 1965). Reykjanes samples are particularly interesting because they were collected on the present Mid Atlantic Ridge crest at a recently emerged segment of the ridge on the southwestemmost point on Iceland. Pillow basalt with some interpillow hyaloclastite is exposed at this location. These deposits appear much less altered and are probably younger (early post glacial) than the deposits along the shore near Reykjavik (Briinnes, Kopavogur, Arnarnesvogur). All of these seawater-erupted glasses are now exposed to meteoric water. Periodic sea level changes and groundwater outflow may have led to a complex alteration history. 26 1.2.2 DREDGE SAMPLES Fourteen samples of pillow rim material and hyaloclastite dredged from the deep sea were included in this study (Table 4). These samples were chosen fran the dredge collection of the National Museum of Natural History, Smithsonian Institution (see Byers, Jercinovic, and Ewing 1986). The samples were collected from the surface of the sea floor and, thus, were exposed to the open ocean environment throughout their histories. Ages of dredge samples are difficult to obtain but can be constrained based on location relative to active spreading ridges, age-correlated magnetic anomalies, or dated basalts in drill cores. In a few cases, microfossil evidence helps to constrain ages somewhat. 1.2.3 DRILL CERE SAMPLES Eight glassy samples from oceanic basement deposits were also included (Table 5). These samples were chosen from the core library of the Ocean Drilling Program (formerly the Deep Sea Drilling Project). They were chosen based on their well known ages, age distribution, and glassy nature. For each hole, information is also available on the overlying sediment cover. Although these samples have been beneath the sea thoughout their histories, it is not known if seawater has had access to the samples for that 27 Tabl* 4. Oatfi (ta drtdia aaaplaa. location pta (nvlr •Marks laykjana» rldaa ti.***! 24.59*u Utm 113521-27 .1 to 1.5 ay faabod Ullon fraaaant with .3 ca thick (laat rfa. 113521-49 taabcd Kilo» ria, 1 ca thick flaaa 59.75*» 2*.I7*W USMN 113531-4 «1.5 mf Saabad Hifhly altarad hyalaclaatita 52.47*» 34.94"V HIM 113*27-25 <30 ay laaoad »aarly unaltarad flaaa •id Atlantic lidft 45*31'» I»*34'« Ul"« 11J4I7-3 9.5 to 11.0 ay toabad Coaalataty altorad hyaioclaatita 23*01'» 45*02'W Ufa» 113715 «1 mt laabod Slightly altarad 1 ca thick tlaaa frafatnt 9*3*'» 40*39'W USNN 11130* <30 m» laabod »llgntly aitarad pfllw ria gltn 0.93*» 2».37*V 110728-448 m laabad Nyaloclaatita braccia Calapafoa lidta Araa Gatapaaaa Ifdft USM 113152-17 <30 t faabad Glaaay FITI aaaalt 2.18°J 101.94** UfNM 113434-3 m lotbad *1 How ria tlaaa • UtNM 113434-15 m taabad • i 11 on ria tlaaa 3.20*1 «5.04*W USM 113249-705 <2 m laabad »illow rfa tlaaa 9.16*» 105.24°V U»H 1U045-3A 1.4 ay faabad • itlow ria tlaaa Sorda (idfa Araa 44.42*» 129.92*W US«M 111238-1 <30 ay iaabad 1 ca thick tlaaa chip with hyaloclaitita brtccfa 28 Table 3. Oeep ••• drill car* soaplee. location taaple1 «•• Imltormtnt taaerke teykjanea t Idle 62°17'« 25°57'W 00» 49 409 28- 1 2. 4 •y' lub-ioobtd »ateaonite breccia 45°31'« 29°29'W 00» 49 410* 2- 1 9. 5 •v2 !t*>e»abed Kyaljclaatlta breccia 00» 49 410» 2- 2 • tub-aeabod Claaay eetvaae Old ttttntU »löt» 3S*J8'W X» 37 3328 0-2 J. 5 •yJ lub-aeab«d Chalk Kith chip of black «taaa rAMut area and p«la|on 'saaoie maabera refer to, reapect 2 TECHNIQUES The sample characterization scheme is shown in Figure 3. Specimen preparation techniques are extremely important as alteration products must be preserved. 2.1 THIN SECTION PREPARATION Thin sections were prepared from representative chips of hyaloclastites and of rim material (section cut perpendicular to pillow surface) of pillow basalts. Chips or slices were initially impregnated with epoxy (under vacuum) then cut or ground to obtain a mounting surface. "Dry" preparation techniques were used throughout, with kerosene as a lubricant and 2-propanol, and acetone, for cleaning. Sic grit and paper were used for grinding, and diamond paste (down to 1/4 micron) for polishing. In order to maximize preservation of weathering products, samples were re-impregnated with epoxy between each 30 Summory of anolyticol strategy Arrows • techniques Boxes • physical or onolyiicol results nanometer scale mineralogy and Chemistry ol surface layers bulk maior element chemistry of omorpnous. alkaline, or hydrous techniques for onolyzing applied amorphous, alkaline, or phases and maior hydrous pnoset techniques element grodients in glass ond poiogonite rnaior element chemistry of crystalline phases texture, mineralogy, limited chemistry elements to be included in quantitative analysis Figure 3. Summary of characterization scheme. 31 step in preparation (except the final polish). Epoxy was cured at roan temperature to prevent devolatilization. 2.2 SCANNINCf BffiTfQW MICROSCOPY Specimens were prepared for analysis by mounting representative chips onto stubs with glue. Microfossils picked from some samples were attached to silver tape, which had previously been bonded to the specimen stub with carbon paint. Samples are coated with carbon and gold-palladium by vacuum evaporation. Conductivity is insured by applying carbon paint across the glue joint. Specimens were analyzed using a Hitachi S-450 Scanning Electron Microscope, epuipped with a Tracor-Northern, TN-2000 energy dispersive x-ray analysis system. 2.3 X-RAY DIFFRACTION Clay minerals were separated by lightly crushing samples with mortar and pestle and disaggregating ultrasonically. Suspended fines were pipetted off and placed on microscope slides and evaporated. This procedure orients the clays for determination of basal (001) spacings. Three such slides were prepared for each sample, one for treatment with ethylene glycol, one heated to 550°C, and one run under ambient conditions. Packed powders were used to determine (060) spacings. Identifications were based on methods outlined in Brown (1961) and Carroll (1970). Zeolites and other authigenic minerals can 32 be identified from whole-rock powders or from size-separates obtained by sieving. Samples were run on a SONTAG PAEW automated diffractometer system. 2.4 TTprfpnw MTCROFROffi AfflLY-fI? Analyses were performed on a fully automated, five-spectrometer JEOL 733 Supexprobe. The Sandia TASK 8 operating system (Chambers 1985) controls data aquisition an the Superprobe and implements corrections for differential matrix effects by the method of Bence and Albee (1968). A major consideration in the electron microprobe analysis of glass (and, by extension, palagonite and zeolites) is the behavior of alkalis, particularly sodium, when exposed to the electron beam (Melson et al. 1977; Jarosewich, Parkes, and Wiggins 1979; Nielsen and Sigurd' son 1981; Strope 1984; Walker and Hewitt 1986). Efforts to minimize the problem of alkali diffusion have included broadening the beam (Natland and Melson 1980), beam rastering (Strope 1984), empirical corrections, and cryogenic techniques (Nielsen and Sigurdsson 1981). Counts (beam current normalized) vs. time relationships were determined for sodium and potassium on a number of points an glasses, palagonites and zeolites using a 20 nA, 2 micron diameter beam. In each case, sodium exhibited a greater relative loss in counts/time than potassium. Up to 59% loss was seen for sodium in glass over the time of a typical ten element analysis. Defocussing the beam to a diameter of 40 microns resulted in no discernable diffusion over the period of analysis of glasses and zeolites. 33 Low current density is clearly an important factor in lowering the rate of alkali diffusion. For palagonite analysis, only a slightly defocussed beam (5 microns in diameter) could be used, necessitating other procedures to obviate elemental diffusion. As suggested by Velde (1984) for clay analysis, the beam current is lowered to 2 nA. This appears to eliminate diffusion in palagonites, glasses, and feldspar standards over a period of several minutes. During calibration and analysis sodium is analyzed first on the appropriate spectrometer. Additionally, the calibration and analysis times for sodium were lowered (40 second count times were used for most elements, 20 seconds was used for sodium). Although this results in lower statistical accuracy for individual points, averaging over many points appears to give reproducible analyses. Another potential difficulty in the analysis of hydrous phases (palagonites, zeolites, clays) is devolatilization. Count rates were not observed to increase for any major elements (silicon and iron were specifically monitored) over the time of a ten element analysis. Neither do analytical totals increase for analyses run repeatedly on the same point on any phase. Local specimen heating due to the electron beam does not raise temperatures above critical dehydraxylation temperatures, otherwise the totals would increase. Adsorbed water in palagonite samples has been observed to be lost when placed under vacuum (Allen et al. 1981), thus adsorbed water is not a factor in the microprobe analysis. 34 Glass and mineral standards were used. Glass, palagonite, and clay calibrations were checked against independent standards and a reference glass (USNM 113716, see Jarosewich, Nelen, and Norberg 1980). Zeolite calibrations were checked against feldspar standards and the reference glass. 2.5 Samples were prepared for ultramicrotomy by impregnation with epoxy resin poured onto the surface of the specimens in vacuum (10~2 torr) in order to insure penetration of the resin through the surface layer. Specimens wer« cut perpendicular to the original surface by a Reichert-Jung Ultracut E-Microtame to a thickness of 50 to 80 nm. The analytical electron microscopy (AIM) was completed on a JEOL 2000FX with a Tracor-Northern TN-5500 energy dispersive spectrometer. High resolution bright field (HREF) and dark field imaging (HRDF) and electron microdiffraction methods were employed along with the usual bright field (BF) imaging and selected area diffraction (SAD) techniques. A Gatan low-background, double-tilt cold stage (-150°C) was used as a specimen holder to avoid element loss and reduce electron beam damage during elemental analysis and high resolution examination. The chemical analyses were completed with a beam size of 20 to 30 nm. Standards used for Na, Mg, Al, Si, Ca, K, and Fe for the thin film quantitative analysis were Amelia albite, albite, sphene, jadeite, hornblende, cummingtonite, and tremolite. All analyses were normalized to 100 percent and errors were estimated to be les than 10 percent. One should note that water (not determined) is an important component in the leached layers. 35 RBS0LT8 3.1 3.1.1 GENERAL All glass samples from Iceland included in this study exhibit evidence of palagonitization, in particular, dissolution-pitted glass surfaces with :iated surface layers which pseudonorpnically replace the glass. These surface layers are red, orange, yellow, or brown and range from optically isotropic and non-crystalline (amorphous gel) to somewhat anisotropic (cryptocrystalline). The thickness of rinds varies from sample to sample (averaging less than 10 microns in some samples to more than 100 microns in others) and, to a lijnited degree, within a sample (generally variation within about a factor of 2 from the median value). Fresh glass remains in all Some samples contain cements, most commonly hydrous aluminosilicates such as smectite clays and zeolites. Calcite occurs in several samples, almost exclusively as a late-stage cement in highly altered deposits. Clays are ubiquitous cements, occuring as minor overgrowths in sane samples and abundant cements, filling all available pore space, in others. In all cases, palagonite is the first alteration phase to form. Partial or complete reconstruction to smectite clay is common. Thick palagonite rinds (several tens to hundreds of microns thick) are often associated with 36 zeolites. Zeolites are not seen where palagonite rinds are thin (below 20 microns). With the exception of high temperature situations, primary crystals (olivine, pyroxene, and palagioclase) are essentially unaltered. This is true even where palagcnite has replaced all glass surrounding the crystal. 3.1.2 FRESH WATER ALTERATION Samples collected from inland Pleistocene subglacial volcanoes (Trolladyngia, Mosfell, Middalsfjall, Raudafell, Blafell, and Valafell) have been exposed to fresh water for at least 8000 years, and in sane cases, as long as 700,000 years. The volcanoes erupt into glacial melt water where the deposits of pillow basalt and hyaloclastite remained subaqueously for the duration of the existence of the englacial lake. During interglacial and post-glacial tine, the deposits are exposed to percolating meteoric water. Plio-Pleistocene (.7 - 3.1 m.y.) hyaloclastite and basalt deposits (Eyrarfjall, Tungufell) have also been exposed to fresh water throughout their history. Active nvdrothermal systems persist in some areas, with associated (and localized) increased alteration effects near the centers of these systems. Glass is devitrified and high temperature minerals (e.g. chlorite, epidote) typify these areas. Except in ran» cases, the age and duration of the hydrothermal systems are unknown. For the purposes of this study, samples from high temperature areas were not included. 3.1.2.1 Pleistocene Subalacial Volcanics 3.1.2.1.1 Palagonite Palagonite forms as red, orange, yellow, or brown surface layers on all exposed glass surfaces (Figure 4). Rind thicknesses (apparent values as obtained fran thin section measurements) are given in Table 6. In general, rind thicknesses are less than 40 microns. The physical character of the palagonite rinds is variable, ranging fran vitreous and isotropic to earthy and sonv+iat anisotropic. There is no correlation between rind character and rind thickness. Ten-element microprobe analyses of glasses and associated palagonite rinds are presented in Table 7. Actual gains or losses of elements can be calculated if phase densities and volume relationships between the initial glass and secondary palagonite are known. Palagonite replaces the glass pseudcBcrphically, retaining even the most subtle of textural features of the glass progenitor (see Figure 4). Thus, the glass to palagonite hydrolysis reaction can generally be considered isovolumetric, that is, the volume of palagonite is equal to the initial volume of reacted glass from which it formed. There is no petrographic evidence in any of the sarnies studied for either shrinkage or expansion, though shrinkage cannot be ruled out for sane high iron deep sea dredge samples. The results of mass per unit volume 38 Figure A. Photomicrograph of a palagonitized glass from Icelandic subglacial volcanic ridge (TroUadyngia). Sample Tr.-1 exhibits thin rinds which replace the glass on exposed surfaces. Note that pa I agon i 11 between glass and the plagioclase crystal could not be of lower volume than the amount of dissolved glass or there would be visible gaps. Table 6. Apparent paiagcr-.te rind thicknesses in Icelandic samples Fresh Water Altered Seawater Altered Sample Thickness (m.crons) Sample Thickness (microns) Tr.- 1 5 • 15 Br.-1 15 • 25 N-l 10 • 80 Br.-2 0 • 440 N-7 5 - 40 Ke.-I 15 - 30 Nid. -1 0 - 40 A-1 40 • 60 Nid. •2 0 - 40 Seyk. -1 20 • 80 Nid. •3 5 • 20 Reyk. •2 20 • 80 R-1 15 • 20 R-2 3 - 5 Bl.- • 20 • 30 Bl .- 4 20 - 30 '•2 ..0 - 18 T-3 20 • 20 T-4 5 • 30 V-1 3 - 10 Apparent range of thicknesses measured in thin section. 40 Taklt 7. *««ri|« alcroproo* tnalyttt <»t. S) of •'••• tnd e*rrtfpondt»t P*i*fenii* tint* frta letlwdlc tub|(*e tr. • 1 »• t »•7 •Id.-I Kid. -2 81... ••(• twit. 81... >«ltf«nfrt SIM* »alafanlt. 01 *M •alafonlta eutt »»tatonlta SIOj 49.7 27.2 48.0 38.1 48.0 38.0 47.9 31.5 48.4 36.2 r(Oj 1.88 3.47 1.53 .71 1.57 2.24 2.54 3.55 2.74 5.7 «l|Oj 13.8 20.3 15.4 12.7 15.4 20.4 13.9 17.0 13.5 14.4 »ao1 12.8 21.7 11.1 12.7 11.2 11.2 13.1 12.8 13.5 17.5 MM .22 .02* .14 .02* .11 .00* .21 .13* .23 .22 "90 4.3 .35 8.7 2.28 8.4 1.49 7.0 1.90 4.8 1.02 c*o 11.1 .52 12.4 1.99 12.3 2.45 11.4 2.42 11.5 4.59 2.33 .14* 2.13 .24* 2.09 .15* 2.23 .24 2.51 .10* *2° .33 .01* .13 .10* .13 .14* .30 .12* .36 .07* •»0« .22 .35 .13 .14 .13 .14 .28 .25 .33 .38 t»til 91.7 74.3 99.9 49.0 99.4 74.5 99.1 77.1 99.9 12.2 ni «.-3 • -1 • • 2CIM1* l*2(p i lieu 1) R-Z(p IIOj 48.1 31.2 48.7 42.1 48.2 33.4 48.5 32.4 48.5 33.4 TIO; 2.60 4.84 i.n 2.46 1.87 3.44 1.87 2.»1 1.87 3.95 *l2?3 13.8 17.7 13.3 16.1 15.2 19.5 15.2 21.0 15.2 20.8 1 '•0 13.3 15.3 11.3 11.0 11.8 12.9 11.4 15.0 11.4 14.4 NfiO .22 .00* .19 .11* .19 .01* .18 .01* .18 .03 »90 7.1 .79 7.9 3.84 7.7 .62 7.5 .43 7.S .54 CM 11.7 1.69 12.1 4.47 11.8 2.84 11.8 1.44 11.8 1.68 I.JO 2.31 .08* 2.09 .04* 2.17 .01* 2.15 .03 2.15 .07* (j0 .32 .06* .27 .01* .28 .04* .27 .02* .27 .07* *2°S .26 .29 .21 .11 .25 .21 .25 .30 .25 .14 fatal 99.7 72.0 99.8 80.5 99.5 75.2 99.1 75.9 99.1 74.2 '«(( ft raoartatf a. HO. *!•(•« 4*tact<*n Malt. Tibl* 7. (continued). .•1 II .•t V-1 Clot» •ol»»on(tt Cl... •ilttontt« Clttl »•l«|onttt 11 Oj 48.7 19.6 U.6 31.2 4*.0 38.2 TtOj 1.60 3.25 1.61 2.37 2.63 3.26 *l2°3 U.6 23.3 U.I 18.9 15.5 13.2 11.C 18.2 10.9 18.1 13.2 13.0 »no .18 .O52 .17 .us 17 .03J "»0 8.0 .58 7.9 .87 7.0 2.81 c »o 12.» 1.26 12.7 1.80 10. 7 4.84 • OjO 2.03 .032 2.06 .202 2.51 .•1* .16 .02* .16 .042 39 .24 t 9 .14 .14 .13 .07 30 .06 76.» 98.6 7J.7 98.4 75.J 42 calculations for glass and palagonite fran Pleistocene subglacial volcanoes are given in Figure 5. The following results are evident: 1) Silica is depleted in palagonite relative to parent glass, as are MgO, cao, Ka2O, and K20; 2) TiC>2, AI2O3, and FeO are approximately unchanged relative to parent glass; 3) Alkaline earth elements are not depleted to the extent that alkalis are. 3.1.2.1.2 Cementation If cementation reactions do not occur, then elements depleted in palagonite relative to glass are removed from the system by flowing water. The formation of iron and/or aluminum rich silicate cements in samples where these elements have been retained during palaganitization indicate an open system as the elements required for these cements must have been externally derived. Bus assumes no shrinkage during palagonitization. Clays are abundant as brown, unoriented matrix in samples from Hos fell. Table 8 gives microprobe analyses of clays in samples M-l, M-3, M-6, and M-7. The compositional relationships are illustrated in Figures 6 through 9. Evidence for mixed-layered clay is clear for clays in samples M-3 and M-7 Tr.-1 i 1 i !"1 * * * • KM me em »•«• n» nm nw tut r* » w0 em tut at nm M-1 Mid.-3 s i T": ! i i n ISM itt n* »i» e— »*n o» nm ct fti at M-7 R-1 r-:- i i i nw i>M nw UN r«« *«< arc CM nut m Mid. —1 R-2 4 -—I i I It T T 10**- «•/» JM nti figure 5. Glass normal ized-palagonite composition of Icelandic subglacial volcanics. vetoes plotting below 1 indicate loss of that element to the I_cai aqueous solution as a result of palagonitization. The oxide concentrations have been convertea :o mass/volume using density values from nay and lijima (1968a; 1968C). An asteri&n indicates that the particular element is below the detection limit in palsgonite. Error bars include one Standard deviation in the microprobe analyses ana tne assumes density ranges. 44 r*~* I" • *••*»**« *M^« v4*****KK"*** * •••••••••• R-2 (pillow 2) -I-.: ...... I] t I I * * 4i M> nil iM Hi m* H0 CW «aM Of MM MM MM tun n» »** a* n*t Hgurt 5. Continutd. Table 8. Average microprobe analyses o' clays in Icelanoic subglaciai volcanics (wt. X). M-1 M-3 M-6 M-7 Hid. -1 Mid.-3 R-1 R-2 SiOj 40.0 44.3 36.2 40.3 42.2 36.7 44.0 35.6 TiOj .71 .48 .52 .66 1.55 1.51 .24 2.27 *l2°3 21.3 14.7 12.5 17.4 17.5 15.8 14.5 18.9 1 8.9 6.7 3.95 10.0 12.0 13.7 11.0 12.7 FeO1 MnO .042 .062 .092 .31 .082 2.05 .20 .032 MgO 2.58 7.2 7.2 4.29 3.92 3.36 10.7 2.81 CaO 2.65 6.4 5.02 4.78 3.25 3.28 3.02 2.58 2 2 Na2O .41 1.07 .98 1.05 .55 .24 .08 .17 2 2 .072 K2° .11 .19 .19 .24 .23 .11 .122 P2°5 .15 .19 .11 .16 .14 .25 .12 .33 total /6.9 81.3 66.8 79.2 81.4 77.0 84.3 75.5 Cavons oasec on 22 oxygens Si 6.5£ 6.96 6.£7 6.62 6.72 6.40 6.77 6.19 T i .09 .06 .07 .08 .18 .20 .03 .30 Al!V 1.42 1.04 1 .14 1.38 1.28 1.60 1.23 1.81 AlV: 2.72 2.72 1.65 1 .99 2.01 1.64 1.40 2.06 Fe 1.23 .88 .63 1.37 1.60 2.0' 1 .41 1.85 Hn .0" .01 .01 .04 .01 .30 .03 .00 «g .63 1.68 2.03 1.05 .93 .87 2.46 .73 Ca .47 1.07 1.02 .84 .56 .61 .50 .48 Na .13 .33 .36 .33 .17 .08 .02 .06 K .02 .04 .05 .05 .05 .02 .02 .02 P .02 .03 .02 .02 .02 .04 .02 .05 sun tetrahedral/suin octahedral cation ratios Fe' only 1.71 1.86 1.82 1.77 1.69 1.59 1.50 1.62 Fe3* only 1.85 1.97 1.89 1.93 1.86 1.80 1.63 1.81 Ca- free Fe2* only 1.62 1.62 1.60 1.59 1.58 1.48 1.42 1.53 F.3* only 1.74 1.71 1.66 1.73 1.74 1.67 1.54 1.71 '»11 Fe reported as FeO. 2gt, 0 w detection limit. 46 tfoctohedral tmaetit* Mita 2J0O i 1.30 f» aaaonlu U-1 Clay triectahedrei emacttte o m All Iron o* F«2+ « • All Iron os r«3+ 1.00 0.00 aio 1.00 i.éo 2.00 Na -t- K Mg+(Mn) M-1 Clay I s s All F« as F»2+ octahedral cations l I* m All F> at Ft3» Ffgurt 6. Variation plot» of •Uroprob» analysts of M-1 elayt. Tha ratio IV/vi raftrs to the total tatrahadral to total octahtdral cation ratio. Tht fialds ara dafincd from analyaa* (n uaavar and Pollard (1973) except for ^••lapenit» which ineludt* analyse reported In Kohyama et al. (1973), Helton and Thompson (1973), Scheideggcr and Stale (1977), Sudo (1978), and Seyfried and Ifschoff (1979). dectonaeroi trrwctjU illit» 2.00 •etojr »wrnieulit»" Te- eopcnltt tnoetanMroi amectitt M-3 clay o • All Iron a» * « All Iron as 1.00 0.6C 0.S0 i.oc Tio 2.00 Nc + K Mg+(Mn) / \ M-3 Clay = s Ai! Fi es Fi2* octahedral cations , t. s Al! 't C* r»J- \ A! vl Fe-(Ti) Figurt 7. Variition plott o^ micreprobe »nilytes of M-3 el»yi. The ratio 1V/VI refers to the total tetrahedrtl to totat octahedral cation ratio. The fields are defines tron analyse* in Weaver and Pollard (1973) except for Fe-iaponite which includes analyses reported in Kohya*a et al. (1973), Mellon and Thompson (1973), Scheidegger and Stakes (1977), Sudo (1978), and Seyfried and iisehoff (1979). The diagonal dasheO lines are the least-squares best fit lines. 48 tfioctohcdrel tmactfte nut» 2.00 ' ...... • IV/V I • t •Mponltt i^.g cfay o • All Iron os F*2« • • All Iron as FaSt 0.00 1.00 lie 2.60 No + Mg+(Mn) M-6 Clay |= • All F» os odahidral cations U s All F« os F«3+ Figur* 8. Variation plots of •fcroprob* analysas of M-6 clays. Tht ratfo IV/VI rtftr* to the total tatrahadral to total octahtdral catfon ratio. Tht fields irt defined from analyses in weaver and Pollard (1973) except for Fe-taponit* »hieh includes analyses reported in KoftyeM et al. (1973), Nelaon and Thompson (1973), Scheidegger and Stakes (1977), Sudo (1978), and Seyfried and Bischoff (1979). aioctehwsrei imactiu iUitl 2.00 ,— ; ly vwrnieuliU" i 1.50 % M-7 clay tnoctonrtrgi vrwcth* o a All iron os F«2+ * s All Iron os F«3+ 1.00 0.00 o.ic 1.00 1.50 2.00 No + K Mg+(Mn) A M-7 Clay - = A; c$ octahedral cations .' = Ai CS Ft*(TI) figure 9. Variation plots of nieroprobc analyses of M-7 clays. The ratio IV/VI refcs to :he total tetrahearal to total octahedral cation ratio. The fields «re defined 'rom analyses in Weaver and Pollard (1973) except for fe-saponite which includes analyses reported in Kohyea* et al. (1973), nelson and Thompson (1973), Scheidegger «nd States (1977), Sudo (1978), and Seyfried and Bisehoff (1979). The diagonal dasnea lines «'e the least-sQuares best fit lines. 50 (Figures 7 and 9). In both cases, one caqponent appears to be tricctahedral smectite and the other an alkali-rich, dioctahedral clay. The observed variation in total alkalis is predominately due to variation in the Na content of the clays. X-ray diffraction results for clay from sample M-3 show 7, 10 and 14 angstrom basal (001) reflections, indicating the presence of three clay components. The high-alkali, dioctahedral component indicated for M-3 and M- 6 is most lilcely a Na-rich illite (bramnelite) or muscovite. An illite-group mineral is a better match than muscovite in the diffraction pattern. Alkali variation is much more limited in clays from samples M-l and M-6. These clays could be alkaline Fe-saponite (which typically gives total tetrahedral[IV]/total octahedral[VI] cation ratios between ideal dioctahedral and ideal tricctahedral smectites) or intermixed dicctahedral smectite and kaolinite. The latter interpretation is favored because of the generally high aluminum content of the bulk material. A signific? ^ kaolinite-group clay contribution is suggested for all Mosfell clays based on the presence of a 7 angstrom diffraction maximum in x-ray patterns from M-3 and M-6. Clays are seen as minor cements in samples Mid.-l and Mid.-3 from Middalsfjall. Analyses are given in Table 8 and compositional variations are shown in Figures 10 and 11. The clays in both these samples fall within the range expected for Fe-saponite with sane variation in alkali content. There is a greater proportion of iron in these clays (30 to 80 percent Fe in octahedral sites) than those seen in samples from Mosfell (mostly 10 to 3(> ant Fe in octahedral sites). The clays in samples from Ranch fell (R-l and R-2) are Fe-saponites (Table 8 and Figures 12 and 13). Microprobe analyses show that the clay in ullt» 2.00 1.J0 Mid.-i clay •etit» o m All Iron a* Ft2+ « x All Iron es 1.0C c.oc o.io 1.Ö0 1.5C 2.Ö0 Na + K Mg+(Mn) / Mid. -1 Clay C$ r«2- _ A;, r, c, AI VI Figure 10. Variation plot* of mieroprob» tnalyttt of Mid.-1 clayt. The ratio IV/VI refers to tht total tatrahedral to total octahedral eation ratio. The fields are defined from am lyte* in Weaver and Pollard (1973) except for fe-sapo"'te «mcfi include! analys** reported in Kohyama et al. (1973), Nelson and Thompson (1973), Scheideager and ftakes (1977), Sudo (1978), and Seyfried and Biscnoff ('979;. 52 diociohwJfol anaetit* iWta 2.00 IB "cloy 1.S0 S lftcL-S clay o * All Iran « « • All Iran M 1.00 0.00 1.00 ilo 2.00 No + K Mg+(Mn) Mid.-3 Clay o * All F« at F«2+ oetahtdral cations • All Ft as Ft3+ Fe+(T!) F 2.00 • é • • "etoy v«rmculiU" w' ft aopchH* R-1 day trioctnhedn* m* 0 m All Iron a* Te2* » s All Iron as r,3* 0.00 0.5C 1.00 1.50 2.00 Na + K Mg+(Mn) = s All F« cs A s All ft ei \ \ Fe+(Ti) FiQurt 12. w«riition plots of fitcroprobt •nily§«» of «*1 eliyt. Tht r«tio IV/V1 reftr$ to the total tttrahtdral to total octahtdral cation ratio. Th* field» are defined fro* analytct in Weaver and Pollard (1973) except for Fe-saponite which includes analyse* reported in Kohyame et al. (1973), Kelson and Thompson (1973), Scheidegger and Stakes (1977), Sudo (1978), and Seyfried and BisehoM (1979). 54 tioctttadrol sfntetitt »ita 'day vwmieulitt* ft itptnlU B-2 Clay trioctshttfrai snwditt a m All Iron as » • All Iron as F*3+ 1.00 0.6c oio 1.00 lio 2.60 No + K Mg+(Mn) Ä-2 Clay s AII r« ct odahtdral cations • All Ft os FtJ-t- AIVI Ff0urt 13. Vtrfttfon plots of •icroprobt anslysts of H-2 clays. Tht ratio 1V/VI rtfors to tht total tttrahtdral to total oetthtdrt'. cation ratio. Tito fields trt dtfintd fro» analyst* in Wttvtr and »ollard (1973) txctpt for Ft-taponitt which includts analysts rafwrttd in Kohya«t tt al. (1973), Nalton and Thoapson (1973), SchtidttgcT and Itakts (1977), Sudo (1978), snd Styfritd and liscnoff (1979). sample R-l, frcm a hyaloclastite ash deposit, is more Mg rich than those fro- sample R-2 (pillow basalt). Palagonite in R-l is also more Mg rich than palaganite in R-2 (refer to Table 7). Minor zeolitization occurred in one sanple from Mosfell (M-6), the only such example from the Pleistocene subglacial volcanics studied. The zeolite covers a fracture in the clay matrix of the hyalcclastite forming a layer up to 50 microns thick. Individual crystals are angular and equant, generally less than 30 microns in size. Table 9 and Figure 14 give the compositions cf the zeolites. The composition and crystal form are typical of fresh water Ca- chabazite (see, for example, Hay 1966). 3.1.2.2 Tonaufei: 3.1.2.2.1 Palagonite The Tungufell area is characterized by pillow basalts and assccialted interpillow hyaloclastite. The pillows have centimeter-thick glassy rims. The glass is clear with few microcrystals and is of low vesicularity (<1O volume percent). Thin sections through pillcw rims show that the glass is palagonitized on exposed surfaces and, particularly, in vesicles. As shown ir. Figure 15, this replacement of the glass is apparently isovolumetric. A striking feature is that some vesicles are completely filled with clay while others have thin clay layers which, in turn, are overgrown by zeolites (Figure 15). Zeolitized vesicles appear to be, on average, larger than clay filled vesicles. Zeolites appear also to be associated with thicker 56 Tablt 9. Average microproba analyses of teoMttt from Iceland (wt. X). T-2 N-6 lr.-2 KO.-1 KO.-1 A-1 ehabazite chabazite phiUipsite chabazite ana 1 c i me chabazite sfo2 49.0 46.6 50.2 43.8 49.6 50.7 l 21.2 21.1 19.4 22.3 21.3 21.1 * 2?3 2 F.01 • 04 .082 .13 .18 .15 .062 MgO .58 .79 .052 .81 .18 .16 CaO 10.4 9.5 4.94 11.3 2.18 5.3 Na2O 1.31 1.00 1.81 3.19 7.2 5.7 KjO 1.22 1.43 5.6 1.51 3.39 1.45 •aO .012 .072 .09 .022 .032 .022 total 83.8 80.6 82.2 83.1 84.1 84.5 Cations based on 72 oxygens Si 23.60 23.33 24.84 21.80 24.00 24.16 Al 12.02 12.44 11.28 13.09 12.14 11.83 Ft .02 .03 .05 .08 .06 .02 »g .42 .59 .04 .60 .13 .11 Ca 5.34 5.09 2.62 6.04 1.13 2.72 Na 1.22 .97 1.74 3.08 6.80 5.27 K .75 .91 3.55 .96 2.09 .88 8a .00 .01 .02 .00 .01 .00 Si/ All Ft reported as FeO. 2Bclow detection limit. K MS Zeolites Na Ca+(Mg~Bc Figure U. Plot of exchangeable cations in zeolites in sample M-6 (Mosfell, Iceiana). 58 (IM HM JIIM Ht Ml Ml iM Hl DM tfO Ctt Utåt Bt '-I »ttlCtM • •M Mta^MfcM* • • • v * •: •T 15. Ikttch of a thfn uetlen of Tunguftll, Iceland lamplt T-2. The fiald of via» it approxiuataly S «m. Also shown art the glass normal ized-petagonitt compositions, an exchangeable cation plot for zeolites, and the relationship between apparent petagonfte rind thickness and vesicle size. palagonite rinds. The rird thickness/vesicle size and cement relationships are shown in Figure 15. For zeolitized vesicles, larger vesicles tend to have proportionally thicker palagonite rinds. Table 6 sunmarizes rind thicknesses. Ccnpositians of palagonite rinds were separated into 2 groups: 1) those rinds associated with clay filled vesicles; and 2) those rinds associated with vesicles containing zeolites. Electron microprobe analyses of the two palagonites, along with those of the glass, are presented in Table 10. Figure 15 shows the glass-normalized palagonite analyses for T-2 and Figure 16 shows the analyses for T-3 and T-4. The following relationships are seen: 1) SiC>2, AI2O3, MgO, CaO, and Na2O are depleted in the palagonites relative to parent glass; 2) TiC2, FeC, and K20 are retained during palagonitization; 3) AI2O3 depletion and FeO enrichment are slightly greater in palagonite associated with zeolite containing vesicles compared tc that of clay filled vesicles. Microprobe traverses of palagonite rinds were also completed for sarple T-2, again separating those rinds associated with clay-filled vesicles and those associated with vesicles containing zeolites. The results for FeC, AI2O3, MgO and CaO are illustrated in Figures 17 through 22. The following observations nay be made: 60 t tal t 10. Avtrtfa »Itrtprtb» tna(yttt Cut. I) tf «(«•• and carrmpandlnf paiaganltt rinrft fraa hr«(Mla*t r-2 T-2 (taailit MtMiatM) (c t ar tttacfttad) T-J T-4 eiatt •tlttanitt ttatt Statt •alafanitt Statt Piltfonltt »10, a.3 45.7 41.3 4». 7 41.5 44.» 47.7 37.1 TIOj 1.27 2.*4 I.J7 2.43 1.27 2.»3 1.30 2.7» *l2°3 U.» . 11.0 15.a 11.9 13.7 11.3 15.» 17.1 10.0 11.0 10.0 14.» 10.1 14.7 9.* 15.2 NflO .1* .04* .1» .07* .1» .It .15 .01* »90 l.l 2.14 l.l 1.9S 1.» 2.09 1.7 1.13 c*o 13.2 4.44 13.2 $.7 12.9 t.» 13.1 2.2» ••2o 2.07 .30 2.07 .Ii» 2.04 .22* 2.02 .02* tjO .11 .27 .11 .31 .09 .25 .10 .0»* »2°5 .11 .09 .11 .07 .13 .12 .11 .15 ittti 99.1 «5.3 99.1 (3.9 99.3 •2.9 91.6 77.2 '»II tt rtparttd it rtO. J Atttctien Halt. T-4 S. i Figur» 16. GLast normil i z»a-pil ggoni t» eompodtion of »«mpl»i T-J tnd T-4 from Tungufeil, Iceltnä. Values plotting below 1 indicate lost of that element to the loesi aqueous solution aa a result of palagonitization. The oxide concentrations have been converted to dass/volume using density values from Hay and lijima (1968a; 1968b). An asterisk indicates that the particular element is below the detection Unit in palagonite. Error bars include one standard deviation in the analyses and the assumed density ranges. 62 T—2 (with zeolites) Traverse 1 0 20 40 100 Position (micron») a AI2O3 • r«0 • M«O C«0 T—2 (with zeolites) Traverse 1 •••••••••••••••••i PooHlon (mtoron*) OM • fo • IV/VI Figurt 17. tttutti oi • nieropreb* travers* ef a pelagonite rind In <(«pl* T>2, Tungufell, Iceland. ThU f» Traverae 1 ef 10 which nere done on rind» associated Mith zeoUt» filled vesicles. The location of this traverse is shown in Figure 15. The treverse begins in pelasonite near the gless/pelagonite interfoce end proceeds outward through the rind. The oxide weight percents shown in the upper disgran are recalculated values, normalizing to 100 X totals for the ten element- oxides enalyzed. The lower diagram gives the corresponding number of cstions assuming s 22 oxygen (or 20 oxygen, 4 hydroxyl) unit ell ss well as the resulting total tetrehedral/total octahedral cation ratios. T—2 (with zeolites) Traverse 2 39 30 - 25 • ••••••**•• ••*•• K 20- IS J' o * ° D_0 '«figO i 10 - O • o o • • »« O 20 40 60 80 100 120 140 160 180 PoiKJon (micron«) o A12O3 r~«O • MgO i CoO T—2 (with zeolites) Traverse 2 4.0 - 3.9 - o 3.0 - o 2.9 - •••••••* *** o*'***c 2.0 - " • ' «• «• • B 1.9 - 1.0 - * O.S - 20 40 00 SO 100 120 140 160 ISO Position (rrdoronf) DM • r« • IV/Vl Figure 18. Results of • microprobt trtverst of • ptltgonitc rind in sample T-2, Tungu'e.L, Icctana. This is Traverse 2 of 10 which were done on rinds associated witn zeolite filled vesicles. The traverse begins in pslegonite near the glass/palagonite interface and proceeds outward through the rind. The oxide weight perctnts shown in the upper diagram are recalculated values, normat'zing to 100 X totals for the ten element-oxide» analyzed. The lower diagram gives tn« corresponding number of cations assuming a 22 oxyqen (or 20 oxygen, u hyorcx/i) unit cell as well as the resutting total tetrahedral/total octahedral cat:on ratios. 64 T—2 (with zeolites) Traverse 8 38 so 28- K 20- • • ( f 0 D O B « o a B la- a B B o ic- t t : i al • • O 20 40 0 eo PotHion (mlerona) a AI2O3 F«O • U«O * CoO T—2 (with zeolites) Traverse 8 20 40 •0 PoaHlon (mieront) 0 Al • ro • IV/V» Figur* 19. Rttultt of a Mieroprobt travtrt* of a patagontta Hnd in t amp 11 T-2, Tungufall, tealand. This (a Travtraa 8 of 10 which utrt dona on r 30 • 25 K 20 • • • 0 • • a i is H a a o o 10 - o • • • • • • 0 20 40 60 »0 1C? 120 140 PotHJon (microne) O AI2O3 * r«0 • MgO 4 CoO T—2 (wit* zeolites) Traverse 9 i 20 40 60 90 100 120 PoeKion (microne) O Al • F« • Figure 20. »*»ult» of I microprooe tr«vtr$* of • paligonit* rind in ttmpte T-2, Tungufe.., Ictltnä. This it T-»v«rte 9 of 10 which wtre dont on rinds issociatcd with zeolite filled vesicles. Tht traverse begins in ptisgonite near the gliss/palagoni ce interface and proceeds outward through the ri.id. The exice weight percent» shown ),, the upper diagram are recalculated values, normaiizing to 100 X totals for the ten element-oxides analyzed. The lower diagram gives the corresponding number ai cations assuming a 22 oxygen (or 20 oxygen, k hyoroxyl) unit cell as well as the resulting total tetrahedrai/total octahedral cat-on ratios. 66 T—2 (with clay) Traverse 2 30- 29- M 20- • • * 1 18- i 1 0 o 0 D a 0 a o 10 - * * * A A A A A 8- • • • • * * • • 0 10 20 40 PoaKfon (m(cron«) O AI2O3 VO • UflO 4 CoO T—2 (with clay) Traverse 2 40 tv/Vl Hgurt 21. Raiultt of a « f O 10 30 PoeKlon (rrWeron») O AJ2O3 4 Coo T—2 (with clay) Traverse 4 4.9 4.0 3.9 - 3.0 - 2.9 i 1 0 0 e a 0 0 a • 2.0 -; • • • • • * 0 1.9 - 1.0 - 0.9 - 0.0 - 10 20 30 40 Position (micron*) AJ • F« • IV/Vl figure 22. Rttultt of • microprobt travtrtc of • ptl»gonit« rind in sample T-2, Iceland. This it Traverse 4 of 6 which were done on rinds associated with vesicles filled with unoriented clay. The traverse begins in palagomte near th glass/palagoni te interface and proceeds outward through the nna, The oime weight pcrcents shown in the upper diagram are recalculated values, norma.-z 2) CaO and AI2O3 decrease from the gl&ss/palagonite interface outward through the rind, with corresponding increases in FeO and MgO; 3) The ratio of tetrahedral to octahedral cations, assuming a clay mineral structure, remains constant accross the rinds. Both types of palagonite are dioctahedral, and compositionally resemble nontronite. 3.1.2.2.2 Cementation Clay minerals completely fill some vesicles and fractures and c -snt much of the interpillow hyaloclastite. The days occur as brown, very fine grained (less than 2 microns), unoriented nwflnpq. x-ray diffraction of this clay indicates an interlayered kaolinite-smectite. Microprobe analysis of this clay (Table 11 and Figure 24) is consistent with this interpretation, having a relatively high aluminum content. Variation of the tetrahedral/octahedral cation ratios indicates that, if it is in fact kaolinite/smectite, the smectite would have to be dioctahedral (nontronite or montanorillonite). T —2 Paiagonite Rind Traverses 80 100 PotrtJon (micron») • F«O21 o AI2O3C2 A F«OC2 T—2 Paiagonite Rind Traverses 24 - 23 - • • • • * •t 22 - 4 • • • 4 21 - • • 4 20 - 19 - K 0 • S A 18 - 17 - B A A I 16 -i IS -• • s ° 14 - • ••: 9 13 i 4 * * ° a a Q 0 0 12 - o o o o o o o ° o o o a 11 - 10 - 0 20 40 60 80 100 120 1*0 Position (micron») a A1203Z9 * F«OZ« « M2O3C4 A r«0C4 Figur» 23. Comparison of mteroprobe traverses of the two pilegonire rind types in simple T- 2, Tungufell, Iceland. The upper ditgr»m comperes feO and »^Oj in Traverse ' of reolite filled vesicles (see Figure 17) and Traverse 2 of clay filled ves^c.es (see Figure 21). The lower diagram compares these elements in Traverse S of :eolite filled vesicles (see Figure 19) and Traverse 4 of clay filled vesictes (see Figure 22). Table 11. Average microprobe analyses of clays from Tungufetl, Iceland (wt. X). T-23 T-24 T-3 T-4 SiO2 45.9 45.8 44.3 46.9 2 TiO2 .35 .03 .43 .34 *l2?3 12.6 7.3 11.6 16.8 4.50 7.5 7.3 5.4 MnO .162 .92 .162 .112 NgO 9.6 17.6 9.7 9.5 CaO 4.41 1.56 4.13 3.35 2 2 Na2O .29 .08 .26 .U .33 .36 .30 .36 .08 ,032 .14 .C32 total 78.2 81.2 78.3 82.9 Cations based on 22 oxygen» Si 7.31 7.21 7.19 7.03 Ti 04 .00 .05 .04 IV Al a69 .79 .81 .97 AlVI 1.67 .55 1.40 2.00 Fe 60 .99 .99 .66 Mn 02 .12 .02 .01 «9 2.29 4.13 2.35 2.11 Ca 75 .26 .72 .54 la 09 .02 .08 .04 K 07 .07 .06 ,07 P 01 .CO .02 .00 sum tetrahedral/sum octahedral cation ratios F.2' only 1.73 1.38 1.66 1.65 F.3' only 1.80 1.46 1.76 1.72 Ca-free fe2* only 1.58 1.34 1.52 1.55 Fe'* only 1.64 1.42 1.62 1.61 1All Fe reported tt FtO. 2lelow detection limit. *Unorianted cement. ^Zeolite associated. T-.J clay No Mg*(Mn) anoc\attd) i A\ \ o a AJ ft si oe(aA*etrol catvoru / A ' \ Figur» 24. Compariton of mieroprobe analyses of T-2 clayi, Tunguftll, Iceland. The two types of clay are shown in the thin-soction »k*rch (field of vitw iporoj:maieiy 5 im). The ratio IV/VI refers to the total tetrahedrti to total octahedral cation ratio. The feilds are defined from analyses in weaver and Pellard (1973) excect for fe-saponite which includes analyses reported in Kohyame et al. (1973), Meiso" and T»ompson (1973), Suheidegger and Stakes (1977), Sudo (1978), and Seyf Clay also occurs as thin cementing layers between palaganite rinds and zeolites in sane vesicles. In these instances it is greenish and oriented perpendicular to vesicle walls, forming layers generally less than 30 microns thick (Figure 24). Microprobe analysis of this clay (Table 11 and Figure 24) reveals that it is Mg-saponite. Zeolites have formed in sane vesicles, associated with palagonite rinds which are somewhat thicker than rinds associated with clay-filled vesicles. These zeolites are blccky, forming masses of crystals which completely fill many vesicles (Figure 15). Individual crystals are from 50 to 350 microns in size. Electron microprobe analysis of these zeolites indicates that they are calcium rich, most likely chabazite (Table 9, Figure 15). An interesting feature is that the more potassium-rich zeolites are only found in vesicles near to the surface of the pillow (the outermost 7 mm, Figure 27). 3.1.3 SEÄWATER ÄIHERATION 3.1.3.1 Samples from the shore near Reykjavik (Brimnes, Kopavogur, Arnarnesvogur) and from Reykjanes Point were all exposed to seawater for some, if not nearly all of their alteration history. The deposits at Kcpagogur and Arnarnesvogur are hyalcclastite breccias. Ihe deposits at Brimnes and Reykjanes are fig 25 ill it» 2.00 - "cley vwmiculita" 1.50 - T-3 clay B m All Iron at • • All Iron oi 1.00 0.00 O.SC 1.00 1.50 2.00 No + K Mg+(Mn) T-3 Clay A (unorirnttd cement) \ All r# c$ octahedral catior^ / \ AII '* es • \ AI VI Fe+(Ti) Figure 25. Variation plot* of micrrprebe analyses of T-3 e'.ays. The ratio 1V/V1 refers to the total tetrthedrii to totat octahedral cation ratio. The fields are oe^nec from analyses in Weaver and Pollard (1973) except for Fe-saponite whicn includes analyses reported in Kohyama et al. (1973), Nelson md Thompson (1973), Scheidesger and Stakes (1977), Sudo (1978), and Seyfried and Bisehoff (1979). 74 Aoetshadrsi smectite Mfta 2.00- "cfcy irnikwllu* 1.50 tFe~Mpeftfts T-4 otay o • All iron a* Fe2+ * B All Iran a FaS+ 1.00 0.60 oio 1.00 2.00 No + K Mg+(Mn) T-4 Clay (unorimitd ctmtnt) octahtdral cation» Fe+(Ti) Mgurt 26. Variation plots of mieroprobt analytt* of T-4 el«yt. The ratio IV/V1 rtf»r$ to tha total tatrahadral to total octahadral eation ratio. Tha fialds art dtfintd fro* analyaet in Waavar and Pollard (1973) axcept for Ft-iaponita which includes analyses reported in Kohyaoa et al. (1973), Be I son and Thompson (1973), Schoidegger and Stakes (1977), Judo (197S), and Seyfried snd lischoff (1979). T-2 zeolites I 20-1 I I 1 5 i -5 2C Dtpth From Pillow Surfoc» (mm) Figure 27. Potiinium content of zeolites (from microproDe «n«lysis) vs. depth from tie pillow surface in sampie T-2, Tungufeil, Iceland. The potassium content is g-vc in percent of exchangcaPle cations (Na, K, Ca, Hg, Ba). The error bars -ectsr: 1-sijma standard deviations of zeolite analyses measurec in a pe-vc..a- ves c.e. 76 dominantly pillow sequences with sane interpillow hyaloclastite. All samples are Pleistccene-interglacial or early post glacial. 3.1.3.1.1 Palagonite As with palagcnite in fresh water altered samples, the palagonite in the seawater altered samples is red, yellow, or orange, optically isotropic (amorphous) or slightly anisotropic. The palagonite forms rinds on all exposed glass surfaces, pseudanorphically replacing the glass from the outer surfaces of grains or pillow rims toward the interior. Pålagor i te in seawater altered samples, like that in fresh water altered samples, replaces the glass iscvolumetrically. In nearly all cases, rind thicknesses are below 100 microns, mostly 15 to 80 microns (Table 6). Vesicle size/rind thickness relationships were investigated in samples from Reykjanes (Figure 28). Independent of vesicle size, 90% of all palagonite thicknesses fall into a narrow region of 18 microns plus or minus a factor of two. Ten element microprobe analyses of glasses and corresponding palagonite rinds are summarized in Table 12. Two separate types of the Kopavogur pal^gonite (sample Kb.-l) were analyzed based on thin section observations. The layer next to the glass is red and is clearly distinguised from the outer, yellow portion (Figure 29). The sample from Kopavogur is the only glass sample studied which exhibits this double zonation. This color difference may result from a change in the oxidation state of iron during palagonitization. In sample Reyk.-2c from the Reykjanes area, interpillow-hyaloclastite o Reyk- 1 o g 50- •^ IA • 0) c «o- » « * u * •O 30- a a» a a • a • a • • a a a • • a * • a a a a a a a a* a* ta m a a a* * a • • mm • a a a a a a a* * < ,0- * • • • a Apparent vesicle size microns) Figure 28. Relationship between epptrent ptlegonite rind thickness and vesicle size pillow rim from seawater altered sample Reyk.-1, Reykjanes, Iceland. Ttbit 12. »v«re«t i•IcroO'OD» analvttt (at. X) 0' tltlt trM COrrt>DO"C!'"g salt »or>tt rtna» friM ttaaaitr altarte fiyaleelattlttt and pi 1 low basal: rimt 'n Iciiane. »r .-1 tr •.-2 [«.-1(r»d-fnntr) (0.-t(y«Uo»-outl-) Clltt •ttttonUt Clttl »tiltonitt Gltll t'tlttomtt Clttl »alagonitt tfo2 48 37, 4 46.6 35.2 46.4 37.1 48.4 39.2 TIOj 2.25 3,.34 2.17 3.02 2.60 4.30 2.60 4.46 4I2°3 13 .7 16..9 13.8 14.4 13.8 9.4 13.8 8.3 13 .1 16..3 13.0 17.1 13.7 16.9 13.6 15.4 •nO .20 .03* .21 .02* .23 .OS2 .23 .15J •to 7.2 1..11 7.1 1.50 7.2 2.05 7.2 4.76 CaO 12 .1 2..22 12.1 3.22 11.8 6.6 11.8 5.3 »ijO 2.30 .28 2.26 .15* 2.11 .89 2.11 .62 JjO .25 ,27 .26 .25 .18 .62 .18 .87 .20 *4 .23 .26 .30 .08 .30 .01 total 100 80. 3 99.8 75.1 100.3 78.0 99.9 79.1 J »• »tyt . -1 •tyk.- 2c(p<:le> 'iai •tvt-i!c{lP») Claai * tor.itt Cittl Clatt ttlager • Clltt Pti tgoni11 sio: 48.5 41 49.3 4C.3 49.7 37.0 49.8 43.4 TIOj 2.5! 4 '.42 1.93 i.at 1.93 2.66 AljC, 13.7 9 < 14.2 8.9 14.2 15.' 14.2 U.3 'to' 'It 12 T 12.1 17.6 12.2 16.4 12.2 15.6 »rC .24 '.092 .19 .062 .17 .0 ? .17 .00J •»0 7.! 1.30 6.9 6.2 6.9 2.63 6.9 3.91 CaC 11.9 7 .5 11.2 .34 11.3 .98 11.3 .69 ntjC 2.16 .34 2.21 2.44 2.21 2.10 2.21 1.79 «2o .19 .18 .30 1.11 .30 1.0C .30 1.06 '2°5 .28 .06 .19 .15 .27 .10 .27 .062 to:»i 100.2 77 .5 96.6 76.5 99.2 80.4 99.3 80.6 'Ml ft rtMrtttf t» '*O. 'Ition atitcKen i laf t. '!n(trpi I Ion «r»ioc:»n I tt. Ko.-' (r»C - innay, I,-, Ml n» ii<«i /» »M att et ftt m tut iw JUM /•» »x *>< (•< »•" "t Figurt 29. Glatt normalized-paiagonitt compoaitiont of ttawater altered tanplc Ko.-I, Kopavogur, Iceland. The thin-teetion aketch (field of view approximately 2.5 mm) ilLuatrate» the two types of palagonite in this sample: 1) red, occurring next to the glass; and 2) yellow which occurs outside of the red palagonite. Values plotting below 1 indicate lots of that element to the local aqueous solution as a result of palagonitization. The oxide concentration! have been converted to matt/voluaw using density values from Hay and lijima (1966a; 1968b}. An astensK indicates that the particular clement is below the detection limit in palegomte. Error bars include one standard deviation in the microprooe analyses and the assumed density ranges. 80 palagonite was analyzed separately from palagonite occurring on the pillow rim. Figure 30 shows the glass-normalized mass per unit volume relationships for each palagonite. The general results are as follows: 1) SiO2, MnO, MgO, CaO, and Na2o are depleted in palagonite relative to glass in all samples; 2) AI2O3 is depleted in palagonite rinds in Kb.-l, A-l, Reyk.-l, and Rpyk.-2c (interpillow hyaloclastite). The Reyk.-2c pillow rim palagonite is significantly higher in AI2O3 compared to the palagonite of the hyaloclastite from the same sazcple; 3) K2O is enriched in the palagonite rinds relative to glass; 4) TIO2 and FeO are retained in the palagonite rinds of all samples; 5) The outermost palagonite (yellow) is more MgO rich than the red, innermost (nearest glass}- palagonite in sample (Ko.-l). Microprobe traverses of palagonite rinds were completed on samples Ko.- 1, A-l, Br.-2, and Reyk.-l. The rasults for FeO, AI2O3, CaO and MgO and presented in Figures 31 through 40. Ihe following generalizations can be made: 81 I- T f — Reyk .-2C Br.-2 (p«*. rtm) I .... flft —-i- -J-i- T r- I J---I- ""T « i I I I Sit'- I i"; I * * MM MM imm ft» ut *f» n* »•» a* /•* DM ««• CM *•» Of »M 1«' Rayk.-2C A-1 (W*) I i T i i T ' T « I 1 I i I I I h I I h I * * MM fUt iB« /* *M KlC CM »Wt B( KM Flgura 30. Slaaa normalfzad-palagonfta composition of ttawttcr altorad sampla* from Ictland. Valuaa plotting baloM 1 fndieat» loas of that atamant to tha local aquaous solution as a rasult of palagonitiiation. Tha oxida concentrations hava baan convartad to aaas/voluaw using dtnsity valua» from Hay and Hjima (1968a; 1968b). An ostariak Indicatas that tha particular alemtnt 30- 29- K 20- 1ft 10 a I ? 9 ° o o o • o «•:.• A * * * A 0 20 40 60 60 a AI203 • r«o « M«O « c«o Br. —2 Traverse 1 4.0 0.0 Ffgurt 31. Rtsulu of • nicroprobt travarsa of a palagenUa find in lampta »r.-2, iHmnai, tcaland. This )• Travarta 1 of 6. Tha travarta bagint 20 30 40 P*aftl«*aftl n (mieronana) O M2O3 F«O M«O A C«0 Br.—2 Traverse 2 Figur» 32. Raaults of a nUrepreb* travaraa of a palagonftt rind in aanpla lr.-2, Irinnat, lsaland. Thia O AI203 • F«O • MgO * C«O Ko. —1 Traverse 2 : • •••*•• • •••••• Ffgurt 33. Rtautti of • mieroprobt trtvtrtt of a palagonitt rind In tanplt K0.-1, Kopavogur, Icatand. Thfa 3O- as 20 * • B • o a o a • 10 O o o o ° A • i s t A • A * 20 40 00 O M203 • r«O • M«M«O A C«O Ko. —1 Traverse 4 0.0 Figur* 34. 'Resulta ef a laicroprobe travtrta of a palagonite rind in sample Ko.-1, Kopavegur, Iceland. Thfa i» Traverse 4 ef 6. See Figure 31 for the location of this traverse. The traverse begins in palagonite near the glass/pelagonite interface and proceeds outward through the rind. The oxide weight percents shown in the upper diagram are recalculated values, normalizing to 100 X totals for the ten element-oxides analyzed. The lower diagram gives the corresponding number of cations assuming a 22 oxygen (or 20 oxygen, 4 hydroxyl) unit call as well as the resulting total tetrahedrel/toul octahedral cation ratios. 86 Ko. —1 Traverse 5 30- 29- 20- • • * O D « 18 i I O • • * * ° o o 4 0 a a O 0 o 0 a 10- 4 4 o a O a 4 * 4 44 4 A 4 • 4 * 5- • t 4 4 • • A * • « • • 4 4 4 0 20 40 60 60 100 Pttettlon (mtorone) O AI309 • F«O • M«O 4 CoO Ko. —1 Traverse 5 4.0 3.5- 3.0- I 2.5 O g o • ° * • • • a 2.0- 1< s :• ': 1 •» • • • 1.0 0.9 0.0 20 40 60 00 100 (ml §> e AJ *T (v/Vi Mgurt 35. »t«ultt of a *icroprobt travaraa of a palagonitt Hnd (n taiiplo Ko.-1, Kopavogur, lealand. Tbfi i» Travarta S of 6. Tha travaraa bag in» In palagonitt naar tha glaat/palagonfta tntarfaca and proeaada outward through tha rind. Tha oxid» uaight pareant» ahown fn tha upptr diagram ara raeaieutatad valuai, normalizing to 100 X totala for tha tan olamant-oxidaa analyztd. Tha lowar diagram givas tha corresponding numbar of cations assuming a 22 oxygan (or 20 oxygen, 4 hydroxyl) unit call at Mil as tha resulting total tetrahedral/total octahedral cation ratios. Ko.—1 Traverse 6 I» 20- CoO Ko. —1 Traverse 6 4.0 3.5 - 3.0 - 2.5 - • • • 8 2.0 H i f • • S • e fl gi i.s o a 1.0 - 0.9- 0.0 20 40 60 80 (mloron») ^ o Al • iv/Vi Ffgurt 36. Rt»ult* of • nieroprobt trtv«r»t of • ptltgonitc rind in »«mpi« Ko.-1, Kocevog,- Ictland. This i* Trtvtrx 6 of 6. The trtvtr»» begins in ptligonite near zr.t glits/palogonite interface and proceeds outward through the rind. The ox^ae weight percents shown in the upper diagram are recalculated values, norm»,i i: ng to 100 X total* for the ten element-ox i des analyzed. The lower diagram gives m corresponding number of cations assuming a 22 oxygen (or 20 oxygen, 4 nyoroxyi) unit cell as well as the resulting total tetrsnedrai/total octanedrai cation ratios. 88 A—1 Traverse 1 rvpvwn ^nwrvni/ O AI2O3 * r«O • M«O 4 A—1 Traverse 1 4.0 Mgurt 37. ftttults of a mtereprobt travaraa of a palagonita rind in ia«plt A-1, Arnarntivogur, lealand. Thla fa Tray*rat 1 of 6. Th« tnvrt» bcgina in palagonftt ntar tht glaaa/palagonit* inttrfaet and procoedi outward through the rind. Tha oxida -tight pcreanta ahown in th« uppar diagram art rtcalculatad valuta, normalfting to 100 X totaia tor tht xtn ai«Mnt-oxido* analyzad. Th* loMtr diagra* givaa tha eorraaponding nuaibor of eationa aaauaHng a 22 oxygtn (or 20 oxygtr,, 4 hydroxyl) unit call aa wall at tha raaulting total tatrah«dral/total octahedral cation ratfoa. 89 A—1 Traverse 4 30- 2B- * 20- • • i 18- l a o a a 10- a 0 a a D 1 S- 4 4 A * 4 4 4 4 4 —f— t 20 40 PoeKion (miorons) a M3O3 • F.O • M«O 4 CoO A—1 Traverse 4 Figur» 38. Results of • mieroprobe trtvtrt» of • paltgonfte rind in simplt A-1, Arnarnttvogur, Ictltnd. Th 30- 28- 20- i 18 • 0 ( i 10- o 5- i * t t 0 2 4 8 8 10 12 14 16 18 20 oaKtofl (mierena) O AI2O3 i «C • MgO é OoO Reyk.—1 Traverse 2 2 4 0 8 10 12 14 10 18 20 0 Al Mgurt 39. Rttults of a mieroprobt travtrta of a palagonitt rtnd in tamplt laykjantt, letland. Thfi fa Travtrat 2 of 6. Tht travcrta bagina in palagonfta naar ttia glaai/palagonitt (ntarfaca and procatda outward through tha rind. Tha ox i dt Might parcanta ahown fn tha uppar diagram arc racalculatad valuaa, normalizing to 100 X totala for tha tan atamant-oxidat analyzad. Tha Lowtr diagram givaa tha corraaponding numtoar of eation» aaauming a 22 oxygan (or 20 oxygan, 4 hydroxyl) uni; call aa wall aa tha ratulting total ratrahadral/total oetahadral cation ratioa. 91 Reyk.—1 Traverse 4 0 4 8 12 16 20 24 26 Poattlofl (micron») O AI2O3 • Ä0 • M«O * C«O Reyk. —1 Traverse 4 4.0- 3.8- 3.0- * 2.5- • 2.0- • • • 9 0 a 6 1.8- 0 1.0- 0.5- 4 • 12 1» 20 24 26 a M '•""E (lf-*T'>IV/Vl figur» 40. Raaulta of a •fcroprcb» travaraa of a palagonita rind in aampl» *»yk.-1, Mykjanat, I c» I and. ThU Is T'svaraa 4 of 6. Th» travtrta begins in pa I agon i i» naar tha glass/palagonit» Intarfaea and procaads outward through tha rind. Tha ox id» wafght oareants shown In tha uppar diagram ara raealeulatad valuas, nomatiiing to 100 X totals for tha tan alawanfoxtdas analyzad. Tha lowar diagra» gives tha corresponding nua>ber of cations assuming a 22 oxygen (or 20 oxygon, 4 hydrcxyl) unit cell aa well as the resultins total tetrahedral/total octahedral cation ratios. 92 1) MgO increases from the glass/palagonite interface outward through the rind in samples Ko.-l, A-l, and Reyk.-l, this increase is quite clear in Ko.-l, and is related to the red to yellow color transition; 2) CaO decreases outward from the glass/palagonite interface in samples Kb-1 and Reyk.-l; 3) AI2O3 decreases outward in rinds in Ko.-l and A-l, and increases outward in Reyk.-l; 4) FeO increases outward through the rinds in A-l and decreases outward in Reyk.-l; 5) TiC>2 decreases outward in rinds in Reyk.-l; 6) Tetrahedral to octahedral cation ratios remain constant accross rinds, with values consistently near 2.0 (dioctahedral smectite). Palagcnite in these samples resembles nontronite. 3.1.3.1.2 Cementation Clays are abundant fine grained cements in the Brimnes and Arnarnesvogur samples. At Arnarnesvogur, these clays may be detrital (detrital material has 93 formed elsewhere and been transported to the sample) as fragmented inorganic and organic debris is included in the cement, forming a nearly opaque matrix. Clays are also present but rare in samples from Kopavogur and Reykjanas. At Kopavogur (sample Ko.-l) the clay occurs as orange brown, isolated patches in pore spaces and fractures (Figure 41) and is overgrown by zeolites. At Reykjanes, clays also occur as isolated masses, plugging pore spaces in hyalcclastite and filling a few vesicles in pillow rims. The Reykjanes clays are accompanied by scattered patches of iron oxyhydroxide. Microprobe analysis of clays from Brimnes, Kopavogur, Arnarnesvogur, and Reykjanes are presented in Table 13. Brimnes clay (Figures 42 and 43) appears to be alkaline Fe-saponite. X-ray diffraction analysis of Ko.-l and A-l clays indicates the presence of kaolinite/smectite. Qompositionally, the Ko.-l clay is Fe and Vq rich, with Mg predominating in octahedral sites (Figure 41). Ihis clay is evidently an alkaline Fe-saponite but with less Fe than Brimnes clays. A-l (Arnarnesvogur) clay can be interpreted in the same way as Ko.-l clay, although A-l clay is ccnparitively more aluminum rich (Figure 44). Clay from Reykjanes is very alkaline (Figures 45 and 46) with the ccapositional variation suggesting mixed illite and trioctahedral smectite components. Chabazite is, by far, the most abundant zeolite seen in seawater altered samples from Iceland. Chabazite is seen in samples from Kopavogur and Arnarnesvogur (Figures 47 and 48), occurring as rhombic (nearly cubic) crystals up to 300 microns in size in Ko.-l and up to 100 microns in A-l. Electron microprobe analysis of Ko.-l zeolites (Table 9 and Figure 41) 94 Ko-1 ZtoUtf Ko.-t Clay •'••'I « • M Inn M fel* IM NO Co*(Mg+8o) 0J0 1Ä Mg+(Mn) r.*(Ti) Ftgurt 41. City and ztolftt composition» (mfcroprobt analysts) in samplt Ko.-I, Kopavogur, letland. Tha thin faction skateh (field of vitM approximataly 2.S im) shows tht patrographie rslationsMps corrasponding to tha analysas. Nota that ana lei ma overgrows chabazite. Tablt 13. Average microprobe analyses of clays in icelanaic altered hyaloclastitcs and pillow basalts (wt. X). Htyk.-I fttyk.-2c Br.-1 8r.-2 A-1 $iO2 45.0 42.9 45.1 38.1 43.7 41.5 TiO2 .71 1.52 .82 .47 .55 .41 Al2°3 11.3 15.7 11.0 10.7 10.7 11.3 14.5 11.9 14.6 8.2 9.1 4.86 MnO .002 .022 .22 .152 .25 .12 M9O 6.7 4.91 7.7 7.6 16.8 10.6 CaO .26 .42 1.34 2.12 3.58 4.67 ka2O 2.20 1.74 .92 .48 1.00 1.85 K 1.24 1.26 1.76 .30 1.12 .40 2° 2 P2O« .01 .08 .09 .08 .92 .50 total 81.9 80.4 83.5 68.7 87.7 76.2 Cation» bated on 22 oxygens Si 7.26 6.92 7,18 7.13 6.51 6 .83 Ti .09 .19 .10 .07 .06 .05 AlIV 1.40 1.08 .82 .87 1 .49 1.IS AlVI 2.14 1.92 1.24 1.49 .40 1.01 Fe 1.96 1.61 1.93 1.28 1 . 14 .67 Hn .00 .00 .03 .02 .03 .02 "g 1.62 1.18 1.82 2.13 3.73 2 .61 Ca .05 .07 .23 .43 .57 .82 Na .69 .55 .28 .17 .29 .59 X .26 .26 .36 .19 .21 .08 p .00 .01 .01 .01 .12 .07 turn tttrahtdral/tum octahedral cation ratios Ft2* only 1.58 1.64 1.56 1.60 1.49 1.76 Ft3* only 1.78 1.80 1.75 1.73 1.59 1.84 Ca-frtt Ft2* only 1.57 1.62 1.52 1.53 1.40 1.59 Ft3* only 1.77 1.79 1.71 1.65 1.49 1.65 1All Ft reported at FeO. 2lelow detection Knit. 96 «ioctah«drol vnaetito suit* 2.00 :_*: i Br.-I elay o • Ail Iran at F*2* A s All Iron as Ft3+ 1.00 0.M 2.60 Na + K lio Mg-l-(Mn) .-/ Clay s All F« OS Ft2+ oetaAcdral cation* • All F« at Ft3+ Ft+(T!) Flgun 42. variation plots of nieroprob* analyaa* of lr.-1 clays "day vermiculite" 130 • fl eopenKt Br.-2 elay trioctthadrol srrwclita o • All Iron as * s All Iran os 1.00 0.00 o.io 1.» 1.50 2.60 No + K Mg+(Wn) Br.-2 Clay octahedral cations AIVI Fe+(Ti) Figur* 43. Variition plots of mieroprob* analysts of »r.-2 clays (Brimnes, Iceland). The rstio IV/VI refers to the total tetrahedral to total octahedral cation ratio. The fields are defined from analyses in Uesver and Pol lera (1973) except for f«- saponitc which Includes analyses reported in Kohyama et al. (1973), Mellon ana Thompson (1973), Scheidegger and Stakes (1977), Sude (1978), snd Seyfried and Bischoff (1979). 98 «1oeUh«r*l wrtaeUt* ilrt» 2.00 i . *ctoy varmicuiHa" *4 * 140 Ti aoaonlU Bl A—1 clay tnoctshMial amactfta o • All Iran oa F»2+ * « All Iron sa Pa3+ 1.00 0,60 Toö iJo 2.60 Na + K Mg+(Mn) A-1 Clay octahedral cations Fe+(Ti) Figur» 44. Variation pleta of af croprob* antlytt* of A-1 claya (Arnarntsvogur, Icaland). Tht ratio IV/VI rafara to th» total tatrahadrat to total octahadrat cation ratio, Th» fields »r* ä»Hn*a from analyaaa in yaavar and Pollard (1973) axcapt for fa- aaponita which ineludat anatyaaa raportad In Kohyama at al. (1973), Matton and Thompson (1973), tchaidtggar and Itaka* (1977), Sudo (1973), and Sayfrtad and •iachoff (1979). 99 ileetahadroi wnactit* vm 2.00 i •* »• * 4toy vannieiAa" , ../. x'j_*ij[.:.'_ , 140 Rryk."1 clay IrtoctttMtffoi «MetRa a • All Iron at • » All Iron as FaS-t 1.00 0.00 oio i3ö 2.00 Na + K Mg+(Mn) Reyk.-1 Clay 3 s All r# et Ft2+ octahedral cation* • AM r< ot r«3-t- AIVI Fe+(Ti) Flgurt 45. Variation plots of aieroprob* analyttt of «tyk.-1 eiayt (Raykjant*, letland). Tha ratio IV/VI rafars to tht total tatrahadral to total oetahadral cation ratio. Tht fltlda ara daflnad fron analyaaa In Waavar and Pollard (1973) axctpt for Ft- saponlta which Includat analysat raportad In Kohyama at al. (1973), Mtlaon and Thonpson (1973), fchaldaggar and (takaa (1977), Sudo (1978), and Sayfrfad and Kachoff (1979). 100 aiu Z.00- r ****** * t. - " "*"" 1 •M MrmleuRe" "day 1.50 « «eponKt clay o • AH iron a* r«2+ * • All Iron as F«3+ 0.00 Tio Téo 2.Ö0 Na + Mg+(Mn) Reyk.-2c Clay octahedral cation* Ft+(TI) figur» ib. Variation plot* of •ieroproba analyttt of X«yk.-2 clay* (Reykjantt, Icaland). Tht rario IV/Vt rtftr» ro tht total tttrahadral to total octahtdral cation ratio. Tht field* ar* dafintd fro* analyaat in Wtavtr and Pollard (19T3) axcapt for Ft- •aponit* Mtiieh include* analy*** rtporttd in Kohyama at al. (1973), Nelson and ThMpson (1973), Seheidtgger and Stakes (1977), Sudo (1978), and leyfried and liachoff (1979). The diagonal deched linas ar* the l*a Figure 47. Scanning electron photomicrographs of zeolites in sample Ko.-1, Kopavogur, Iceland: a) chabazite; b) analcime; c) heulandite (?); d) zeolites overgrowing calcite. 102 Figure 48. Scanning electron photomicrograph of a) chabazite and b) calcium-silicate (with high P and some AD mineral (tobermorite?) in sample A-1, Arnarnesvogur, Iceland. 103 indicates that there are both Ca-rich and Na-rich varieties present. Scanning electron microscopy of Ko.-l reveals the presence of chabazite (the Ca-rich zeolite), analcime (the Na-rich member), and a platy variety belonging to the clinoptilolite group, possibly heulandite based on morphology (Figure 47}. Oocgjositionally, the chabazite in Arnarnesvogur sample A-l is quite high in Na, with Na generally exceeding Ca on a molar basis (Figure 49). Phillipsite, a K-rich zeolite is seen as a rare encrustation of prismatic crystals growing on clay surfaces in isolated areas of sample Br.-2 from Brimnes (Figure 50). Microprobe analysis reveals its K-rich character (Table 9, Figure 51). Calcite occurs in samples from Kopavogur and Arnarnesvogur. In Kcpavogur sample Ko.-l, large acicular calcite crystals occur up to 400 microns long, growing on clay surfaces and encrusted by zeolites (Figures 41 and 47). In Arnarnesvogur sample A-l, calcite is seen as a late stage cement overgrowing zeolites in sane pore spaces. Sample A-l also has rare precipitate which SDVEDS analysis shows to be dominantly Ca, Si, and P, with some Al (Figure 48). The proportions of Si and P vary considerably from point to point on this material. Form and qualitative composition suggest that this precipitate is tobermorite. The alteration-paragenetic sequence, based on overgrowth relationships (Figure 41), observed in sample Ko.-l (Kopavogur) is: palagonite-clay- calcite-chahazite-analcime. In sample A-l (Arnarnesvogur), the sequence is: palagonite-clay-chabizite-calcite. 104 K A-1 Zeolites Na Ca+(Mg+Ba) Mgurt 49. Plot of txchangtablt cationt in ztoKttt in tampl* A-1 (Arnarnttvogur, Iceland), 105 Figure SO. Scanning electron photomicrograph of pnillipsite in sample Br.-2, Brimnes, Iceland. 106 K Br. -2 Zeolites Na Ca+(Mg+Ba) figu-e 51. Plot of txchtngciblt cations in ztolitts in ttmplt Br.-2 (Brimnts, Iceland). 107 3.2 Fourteen dredge samples, which are part of the U.S. National Museum of Natural History, Smithsonian Institution dredge collection, were included in this study. These samples are of hyaloclastite and glassy pillow basalt rims which have been exposed to the open ocean essentially throughout their history. Although precise ages are unknown, microfossil evidence or the location of the dredge in relation to age-correlated magnetic anomaly patterns can be used to constrain the ages of sane of these samples. 3.2.1 Falagonite Falagonite rind character is highly variable from sample to sample. In sane cases, it is optically clear, vitreous yellow, isotropic material (e.g., USNM 113715). In other instances, it may be deep red and range from vitreous to earthy with regions that are optically anisotropic, suggesting sane phase seggregation(e.g., USNM 113521-69). Highly palagonitized samples (e.g., USNM 113487-3) have palagonite which is largely recrystallized to smectite clay. The apparent rind thicknesses are listed in Table 14. Rind thickness varies considerably within sane samples, a phenomenon sometimes associated with thinning along fractures with penetration into pillow rim glass. Electon microprobe analyses of the glasses and associated palagonite rinds are presented in Table 15. Glass-normalized palagonite compositions (mass/volume) are shown in Figure 52. Dredge sample palagonites occur in two 108 Table 14. Apparent palagonite rind thicknesses in deep sea samples. Dredge Samples Drill Core Samples Sample Thickness (microns) Samp e Thickness (microns) USNM 113521-27 40 - 60 OOP 49 409 28-1 40 • 60 USNM 113521-69 1 - 102 OOP 49 410A 2-1 20 - 50 USNH 113531-6 10 - 40 ODP 49 410A 2-2 40 • 1103 USNM 113427-25 80 - 220 ODP 49 410A 2-2 250 - 24004 USNM 113487-3 >1200 ODP 37 332B 6-2 20 • 80 USNM 113715 10 - 30 OOP 37 332B 45-1 20 - 80 USNM 111308 10 - 20 OOP 37 334 16-5 >3000 USNM 110728-468 40 - 150 ODP 37 334 17-1 10 - 1003 OOP 37 334 17-1 >1700* USNH 113152-17 203 - 2004 ODP 37 335 10-4 20 - 200 USNM 113434-3 20 - 403 USNM 113434-3 60 - 1504 USNM 113434-15 100 • 300 USNM 113249-705 >100 USNM 114045-3A 40 • 400 USNM 111238-1 5 - 103 Apparent range of thicknesses measured in thin section. General case for this sample although some areas exceed 100 microns in thickness. Interior fractures. Outer surfaces. 109 Table 15. Average mi croprob* analyses (wt. X) of glass and corresponding palagonite rinds in deep sea dredge samples. USNN 113521-27 USNM 113521-69 USNM 113531-6 USNH 113427-25 Glass Palagonitt Class Palagonite Glass Palagonite Glass Palagonite SiO2 50.9 10.0 50.8 11.1 50.1 12.8 49.9 11.2 TiO2 1.60 4.00 1.60 3.93 1.11 2.44 1.27 3.43 M2?3 U.O 14.9 14.1 10.4 14.3 10.7 14.7 15.9 FeO1 12.3 31.5 12.5 27.8 11.8 32.7 10. T 28.9 NnO .21 .082 .18 .072 .19 • 142 .19 .29 NgO 6.9 4.42 6.8 1.91 8.0 3.78 8.2 2.69 CaO 11.3 2.22 11.2 1.51 12.0 1.82 11.7 1.59 N«20 2.09 2.67 2.11 2.56 2.12 1.65 2.30 1.40 K .18 .25 .08 .35 .03 .26 .06 .23 PiO2°e .10 3.40 .14 1.61 .12 1.35 .13 2.77 total 99.7 73.4 99.5 61.2 99.8 67.6 99.2 68.4 USNM 113487-3 USNM 113715 USNM 111308 USNM 110728-468 Glass3 Palagonite Glass Palagonite Glass Palagonite Glass Palagont te SiOj 51.2 27.9 50.4 41.6 51.0 23.4 52.2 55.3 TiO2 1.41 2.55 1.86 1.83 1.80 3.41 1.85 1.79 l 15.8 14.1 15.3 18.2 15.2 18.2 15.3 14.3 * 2?1 3 FeS 8.3 15.6 10.0 13.0 10.2 19.0 9.6 9.5 MnO .13 .012 .19 .022 .16 .052 .18 .062 ago 7.2 3.21 7.7 2.78 7.5 1.59 6.3 9.2 CaO 10.8 .37 10.8 1.57 10.9 1.41 10.5 • .52 Ma2o 2.79 1.36 3.02 1.91 2.99 1.60 2.84 2.15 K2O .58 2.99 .14 2.29 .12 .52 .77 1.90 P->O< .21 .54 .19 .26 .18 1.19 .34 .032 total 98.4 68.6 99.6 83.5 100.0 70.4 99.9 94.8 'All ft reported as FeO. 2Below detection limit. JGlass analysis from OOP 49 410A 2-1. 110 Table 15. (continued). USNM 113152-17 USNM 113434-3 USNM 113434-15 data Palagonite Glatt Palagonita Glatt Palagonite SiO2 47.5 22.5 51.1 43.1 50.4 41.3 TiO2 3.07 3.58 2.80 3.06 2.81 3.29 *l2?3 11.5 3.04 12.8 7.4 12.9 7.2 F.01 16.9 27.5 14.1 20.1 14.2 20.5 MnO .24 .16 .23 .092 .19 .052 NgO 4.89 2.05 5.8 2.14 5.8 2.15 CaO 9.3 1.82 9.8 1.10 9.5 .75 Na2O 2.38 2.63 2.82 2.29 2.89 1.91 K20 .18 .91 .17 2.42 .03 2.70 P2O5 .39 1.47 .27 .32 .24 .18 total 96.4 65.7 99.9 82.0 99.0 80.0 USNM 113249-705 USNM 114045-3A USNM 111238-1 Glatt Palagonite Glatt Palagonite Glatt Palagonite s A12O3 15.5 5.2 17.2 7.6 13.1 7.3 F.S1 8.8 23.0 8.1 20.1 12.9 25.7 MnO .15 .81 .14 .29* .23 .002 MgO 8.3 2.41 9.6 2.93 6.5 4.79 CaO 12.6 .86 12.3 .37 11.0 .42 Na2O 1.86 2.44 2.50 1.70 2.67 1.41 .04 1.59 .04 3.62 .17 2.55 "2° 2 P2O5 .10 .37 .11 ,12 .21 .01 total 100.0 76.2 99.9 82.0 99.3 88.1 , 'All Fe reported at FeO. *telo* detection limit. Ill 10* USNU IIMI7-) USNM 11J71S T • I T T T T c/glos s i I ""i "T I I" I I I i- I, I * I f MM «t w itt» a* MM nu ta*» r* ihé *$» cm* >*n a» USNU ti '*1 n 3 i -i I f I * MM n*t iUM ** If* M* Cmt DM* Of USNU 11MJ4-) USNU HJ4J4-1J I i i ...... £ i,.- — ...... "I" 1 ^ I l"".\ I [ .• äi* c* nut a* MM MM MU ft* 1*4 KfO C* Hut CM HH figura 52. Glasa nor«al Low iron palagonites (Figure 52) are compositionaliy similar to seawater altered samples from Iceland except: 1) K2O enrichments are generally greater than seen in the Icelandic 2) AI2O3 depletions are seen in a number of dredge sample palagonites, generally associated with apparent HO2 losses. These may not be true depletions if the assumed densities are too low or if the layer occupies more volume than the glass precursor. There is no petrographic evidence for volume increase in these samples (e.g., wedging open of fractures). Revising the palagonite densities upward to allow conservation of TiC>2 during palagonitization would eliminate the apparent AI2O3 depletions except for USNM 113487-3, USNM 113434-3 and USNM 113434-15 which would still exhibit A12O3 losses. Note that, in order for titanium to be conserved without layer expansion, densities for the alteration layers would have to be greater than the densities of corresponding glasses. The TiO2 depletions observed in these samples are probably real; 114 3) In all dredge sairple palagcnites, MgO is retained more effectively than CaO, an effect not consistently seen in the Icelandic samples. Similar to low iron dredge sample palagonites, the high iron palagonites (Figure 53) also show greater retention of MgO than CaO. TiC>2 and FeO are not only conserved during hydrolysis of the glass but show significant enrichments above the amount of titanium and iron available from dissolution. TIO2 and FeO are enriched by approximately equal amounts. Some shrinkage during palagonitization may have occurred in these samples, such that Fe and Ti would not neccessarily be enriched beyond the amount made available by glass dissolution. Assuming palagonite densities in the range of iron hydroxides (2.7 to 4.3 g/cro3), the AI2O3 content appears to be constant from glass to palagonite. SiO2 would then have values similar to values for low iron dredge sample palagonites. However, this also results in apparent enrichments in FeO and TiO2 which are even higher than the enrichments indicated with lower densities (approximately double, see Figure 53). Na20 is retained, or perhaps even enriched depending on assumed densities in the palagonite rinds. This is different from the high-Si, low-Fe palagonite dredge samples and seawater altered Iceland samples which exhibit at least partial sodium loss. The ratio of Fe to Na is generally greater than 10 in the high-Fe dredge sample palagonites and less than 10 in low-Fe dredge sample palagonites indicating that there is, in fact, an iron enrichment in the high-Fe palagonites if the apparent high Na content in the same samples is real. The required shrinkage factors for constant Fe during palagonitization would have to be as much as 51% (USNM 113531-6) assuming a palagonite density of 2.04 (Table 16). Lower density for palagonite would result in correspondingly lower shrinkage factors for the constant Fe condition. High mass loss during palagonitization need pologofiiie/glos* potagoniie/glost ni — ; å 0 —• ! |! % : 1 O n o • c 3 a. -> o a <• O c • o •- 3 c •- < S • a. •< c Q. — 3- o «-• • 3 • 2. Us: •3_ TI S ~ 3 . O — -fc -O — -• — O r« fl • *• — • 3 3 • c 5 ? 2 S a m • -I 96 1 — o -» S a » patogonite/glosi O. i» ' '. : • . " a. 3 -> — o in i O 3 i 1 -. S. ti • • M 1 M 1 3 — <• 3 O i "• s l-^H K i • S å • — — < -« Ui n o 7 ~ s i O 7 • - •+ C i -o •« -• a • o- c c a. t HM Ul 116 Table 16. Results for constant F« in the high-Fe, low-Si deep sea dredge pa I .agon i tes. Shrinkage elemental ms» loss (X) total Sample Factor (X)1 Si Ha Mt. X2 USNM 113531-6 51 91 72 67.6 USNH 113427-25 49 92 78 68.4 USNM 113521-27 47 92 50 73.4 USNM 113521-69 38 90 46 61.2 USNM 113152-17 16 71 32 65.7 1 [(Volume glass dissolved)-(voluae palagonito/volume glass dissolved] ' 100 'Analytical total from electron microprobe analysis. 117 not necessarily result in a rind which is thinner than the glass which it replaces. The material may take a more open structure thus having less total mass in the same volume. This would be seen as lower density and should be reflected in lower analytical totals. As can be seen in Table 16, there is no clear relationship between the calculated shrinkage/density factor and total wt.%. Micrcprobe traverses of palagonite rinds were completed on samples USNM 113715 (low-Fe, high-Si) and USNM 113521-69 (high-Fe, low-Si). No discernable compositional gradients were detected (Figures 54 through 60). FeO and AI2O3 are inversely correlated. 3.2.2 Cementation Clay minerals commonly accompany palagonitized glass in dredge samples, although much is of detrital origin (such as in USNM 113531-6 and USNM 110727- 25) and thus are not alteration products. Authigenic clays are seen in USNM 111238-1, USNM 113521-69, USNM 114045-3A, USNM 113434-15, USNM 110728- 168, USNM 113531-6, and USNM 113427-25. Electron microprobe analyses of the clays in samples USNM 111238-1 and USNM 113487-3 are given in Table 17. In both cases, the clay is chemically an iron sapcnite. The matrix material which cements hyaloclastite in USNM 113487-3 is authiginic aluminum-bearing iron oxide. This sample also contains small amounts of phillipsite which occur as a fillings at the cores of sane completely altered glass fragments (Table 18 and Figure 61). 118 USNM 1 13715 Traverse 1 30- 25- 0 O a a 0 0 a M 20 i i 0 •• i is - • a • • 10- 5 - t A A A 0- A A A 0 10 20 30 Poattion (micron») a AJ2OJ • r«o • M«O A Coo USNM 113715 Traverse 1 40 PoeKlon (mierona) o Al • fm • IV/Vi Figura 54. Raaultt of a nicroproba travarta of a palagonit» rind in deep ««a dr*dg« sample USNM 113715 (Low Fao, high S(02). This is Travarta 1 of 6. Tha invert* begin» in palagonitt naar tha glasa/palagonite interface and proceed* acrot» the rind (which occur» in a fracture) to near the glats/palagonite interface on the other aide. The oxide weight percent» shown in the upper diagram are recalculated values, normaliiing to 100 X totalt for the ten element-ox i det analyzed. The lower diagram givet the corrctponding number of cation» •»turning a 22 oxygen (or 20 oxygen, i hydroxyl) unit cell as -ell a» the retulting total tetrehedral/total octahedral cation ratios. 119 USNM 113715 Traverse 2 K 20- 20 40 SO Poettlon (mJcron») AI203 F«O M9O CoO USNM 113715 Traverse 2 4.5 Position (micron*) a AI 4 ro • IV/Vi Ffgur* 55. Rtsults of • mieroprob* trtvtrt* of • ptlagonftt rind in dttp *•» drtdge sample USNM 113715 (low ftO, high SiO2). Thft it Travtrtt 2 of 6. Th« travers* begins in palagonit* near the glass/palagonite interface and proceeds across the rind (which occurs in a fracture) to near the glasi/palagonite interface on the other side. The oxide weight percents shown in the upper diagram ire recalculated values, normalizing to 100 X totals for the ten element-ox i des analyzed. The lower diagram gives the corresponding number of cations assuming a 22 oxygen (or 20 oxygen, 4 hydroxyl) unit cell as well as the resulting total tetrahcdral/total octahedral cation ratios. 120 USNM 113715 Traverse 4 30- a a a 0 25 - O 0 l i a a o 20 - O • 15- • * * * • 10 - 5- • • • • * » • * * A 20 40 60 Poettlon (micron») 0 M2O3 coo USNM 113715 Traverse 4 PoaHlen (micron*) • F« • iv/Vi Figur* 56. Rttultt of a mieroprobt travtrt* of a palagorrit* rind in d«»p «•• dredge (ample USNM 113715 (low F*0, high SiO2). This it Traverse 4 of 6. The traverse begins in palagonit* near the gleis/palagonite interface and proceeds across the rind (which occurs in a fracture) to near the glass/palagonite interface on the other side. The oxide weight percent» shown in the upper diagram are recalculated values, normalizing to 100 X totals for the ten element-ox i des analyzed. The low»; diagram gives the corresponding number of cations assuming a 22 oxygen (or 20 oxygen, 4 hydroxyl) unit cell as well as the resulting total tetrahedral/totat octahedral cation ratios. 121 USNM 113521-69 Traverse 3 4 2 x • • i I I * 0 20 eo 80 Po*Ki»n (micron*) a AJ2O3 • FmO « MgO A C«O USNM 113521-69 Traverse 3 10 80 Poaftlon (mlorona) ^ O M • F« • IV/VI Ffgur» 57. ««»ultt of • »Icroprob» trtvtrtt of • ptlagonitc rind in dttp •#• drtdgt sample USNM 113521-69 (high FtO, low SfO2>- Thi» ii Travart» 3 of 6. Tht travana begins in palagonitt naar tha glas»/palagonfta intarfaca and procaads acrosi tha rind (which oecurt in a fractura) to naar tha gtaat/pelagonita intarfaca on tha othar « 0 20 80 80 100 120 Position (micron») O M2O3 c«o USNM 113521-69 Traverse 4 9.0- 8.0- • • • 7.0- • • • • 6.0- 8.0- 4.0 - a 3.0- a 2.0- a 1i • • • 1.0- 0.0- 20 40 80 80 100 120 PooHlon (micron») O At • r« • IV/VI Figur» 56. Rttult» of a microprob* travtria of a palagonita rind in d»»p ••• drtdga »ample USNM 113521-69 (high F»0, low SiO2>. Thia ft Trävara* 4 of 6. Th« traverse begina in palagonit* near tht glaaa/palagonit* interface and proceed* acron the rind (which occurs in a fracture) to near the gtass/palagonite interface on the other side. The oxide weight pereents shown in the upper diagram ere recalculated values, normalizing to 100 t totals for the ten element-oxide» analyzed. The lower diagram gives the corresponding number of cations assuming a 22 oxygon (or 20 oxygen, 4 hydroxyl} unit cell as well as the resulting total tetrahedrtl/total octahedral cation ratios. 123 USNM 113521-69 Traverse 5 Petition (micron») O AJ2O3 r«O « MoO * c«o USNM 113521-69 Traverse 5 9.0- 8.0- 7.0- S.O - 5.0- 4.0- 3.0 • 0 o a 0 ° 2.0- i • a • • o 1.0 • a 0.0- 20 40 eo Position (micron*) • Fe « IV/Vi Figure 59. Rtsultt of a microprobe traverse of a palagonite rind in deep sea dredge sample USNM 113521-69 (high FeO, low SiO2). This is Traverse 5 of 6. The traverse begins in palagonite near the glass/palagonite interface and proceeds across the rind (which occurs in a fracture) to near the glass/palagonite interface on the other side. The oxide weight percents shown in the upper diagram are recalculated values, normalizing to 100 X totals for the ten element-ox i des analyzed. The lower diagram gives the corresponding number of cations assuming a 22 oxygen (or 20 oxygen, 4 hydroxyl) unit cell a* well as the resulting total tetrahedral/total octahedral cation ratios. 124 Dredge Sample Traverses — SiO2 BO • 90 - K 40 - 30- 20 - o a °° a a Q c 10 i 1 0 0 a a 0 20 40 BO 80 100 120 Petition (mlerena) O USNM 113521-6» #4 • USNM 113719 #4 Figure 60. SiOj vari it i ont tcroaa palagonfta rind» in lanples USNM 113715 (Traverse 4) and USNM 113521-69 (Travaraa 4). Travartaa bag in in palagom'ta naar th« glata/palagonfta intarfaca and proeaad acroaa th» rinda (occurring in fracturai) to naar tha glatt/palagonita inttrfaet on tha othtr tida. Tha oxida waight parcantt »hown ara racalculatad valuas, normalizing to 100 X totals for the ten a tenant-oxidea analyzed. 125 Tabta 17. Avarata »ieroproe» tnalytat af cityt fras doop t« drod«a and drill cort IMPIM (ut. X). UWN 111238-1 Utm 111238-1 Him 113487-3 00» 37 3321 45-I3 00* 37 3321 45-1* no. 41.2 44.7 27.9 37.3 36.3 TIOj .20 .24 2.55 1.92 1.81 *l2?3 3.44 5.» 14.1 12.» 11.5 FoO1 21.0 12.9 IS.» 16.5 IS.3 HnO .012 .06* .012 .102 .132 HtO 10.5 13.1 3.21 2.61 3.42 C tO 1.27 .52 .37 8.7 7.4 *a20 2.11 1.81 1.3» 1.29 .70 tjO .»1 1.22 2.99 .14 .23 »2°S .38 .012 .54 .13 .O52 total 80.7 80.2 »8.» 81.5 7».8 Cat < ont batad wi 22 oiyaont SI 7.1» 7.39 5.73 6.40 ».52 TI .03 .03 .39 .25 .24 AlIV .71 .»1 2.27 1.60 1.4a Al"1 .00 .48 1.14 .94 .95 Fa 3.05 1.78 2.68 2.3» 2.30 Nn .00 .01 .00 .01 .02 »• 2.72 3.23 .98 .6* .92 Ca .24 .09 .08 1.59 1.4] •a .71 .58 .54 .43 .24 C .14 .2» .78 .03 .05 • .0» .00 .09 .02 .01 tua tatrakadral/tia octaJitdral cation ratloa fa2* only 1.41 1.43 1.54 1.90 1.81 fa'* only 1.6* 1.60 1.80 2.22 2.10 Ca-froa »a2* only 1.38 1.43 1.33 1.55 1.51 »o1* only 1.*3 1.SS 1.79 1.80 1.74 'AU Fa roportad at FoO. 2lalM datactfon Holt. 'ironn aatarfal caapaalnf I (war port font af thtrdt. *Ttllo» Mtarlal on thtrd adfat. 126 Table 18. Average microprobe analyses of deep sea zeolites (ut . X). USMM 113487-3 OOP 37 335 10-4 OOP 49 410A 2-1 phi Uipsite phi Ilipsite phi I lipsite SiO2 52.9 53.8 56.3 Al2°3 19.3 19.0 19.0 FeO1 .48 .092 .19 MgO .14 .062 .082 CaO .16 .15 .062 Naz0 6.8 6.1 5.9 K2O 5.3 6.9 6.7 BaO n.3 .032 .012 total 85.1 86.1 88.2 Cations based on 72 oxygens Si 25.24 25.50 25.89 Al 10.86 10.64 10.31 Fe .19 .04 .07 "9 .10 .04 .06 Ca .08 .08 .03 Ma 6.29 5.57. 5.25 K 3.23 4.16 3.91 Ba nt3 .01 .00 2.29 2.39 2.49 1A11 Fe reported as FeO. 2BetOH detection limit. 3Not analyzed. 127 K USNM 113487-3 Zeolites Na Ca- (Mg + Ba) figur* 61. Plot of axchingtibl* citiont in philliptitt in dttp st* dredgt sampl* USNM 113487-3. 128 3.3 CPHi, npRE SAMPLES Eight samples from drill cores of the Ocean Drilling Program (formerly the Deep Sea Drilling Project) were included in this study. The samples, their locations, and ages are listed in Table 5. 3.3.1 Palagonite Table 14 gives the apparent rind thicknesses for deep sea drill core samples. There is no evidence for shrinkage or expansion of the palagonite layer relative to precursor glass in any of the drill core samples. Sample ODP 37 332B 6-2, palagonite occurs as rinds on hyalcclastite grains. The rinds were evidently stripped off of many of the glass shards during transport and deposition of the deposit as many of the grains are free of rinds while other grains have rinds on some exposed surfaces but not on others. The sample is completely cemented by calcite, including abundant calcareous nannofossils (foraminifera are particularly plentiful). Sample ODP 37 332B 45-2, from the same drill site, is a completely altered hyaloclastite. The replaced shard forms are still clearly visible and olivine and plagioclase crystals within the shards are unaltered. The palagonite which replaced the glass is entirely recrystallized to clay. Palagonite in sample ODP 37 334 16-5 is a yellow to yellow-green layer replacing glass on approximately the outer 3 on of a pillow rim. This material varies from completely isotropic (optically amorphous) to 129 recxystallized. Plagioclase crystals completely surrounded by palagonite are unaltered. In ODP 37 334 17-1 the palagonite is of similar character. Penetration of palagonite along fractures has left isolated patches of glass on the outer rim of a pillow basalt. In sample COP 37 335 10-4, palagonite is yellow-orange, generally more orange toward the glass surface. The layer is uneven, from 60 to 200 microns in thickness, and thins along fractures into the glass. The palagonite is apparently recrystallized to verv fine grained (sub-micron grain sizes) clay. ODP 49 409 28-1 palagonite is yellow, forning layers on exposed hyaloclastite surfaces and in fractures. The rinds are of variable character, ranging from isotropic to somewhat recrystallized. Rind thicknesses range from 40 to 60 microns. Palagonite in ODP 49 410A 2-1 (hyaloclastite) has formed as dark red, isotropic layers on glass shard surfaces and along fractures. Rind thicknesses vary from 20 to 50 microns. In some places, the rinds have been stripped away from the glass surface. Rind fragments are seen in the detrital calcite matrix. The other sample from this drill site, ODP 49 410A 2-2, is also a hyaloclastite. Palagonitizatian along fractures is similar to fracture palagonite in ODP 49 410A 2-1. Palagonite preserved on outer surfaces of glass shards is much thicker than the palagonite in fractures (see Table 14). Hicroprobe analyses of glasses and oorresponding palagonite rinds are given in Table 19. Glass normalized palagonite compositions are illustrated on a mass/volume basis in Figure 62. The following conclusions can be made: 130 T06I0 19. Avortoo •icroprob» ono(y*oo (Ht. X) Of-'Oil ond corr-oopondino. polosonit» rindo OO» 49 409 28-1 00» 49 410A 2-1 00» 49 410* 2-2 00» 37 3328 6-2 00» 37 J328 45-11 61 MO »olotanlto Cloot t•oltoonito C1000 »olotonlto 6 lon »olofonito Clooi ••logonit» K Oj 49.8 42.6 51.2 51.5 50.7 44.8 51.0 45.7 s:.o 37.5 TIOJ 1.82 2.14 1.41 2.89 •.44 1.76 1.13 .53 1.13 1.92 "203 13.7 14.0 15.8 15.5 15.7 11.9 14.5 9.5 14.5 12.6 fOO1 12.6 13.1 8.3 18.6 8.2 14.0 10.3 14.3 10.3 16.5 NnO .21 .10 .13 .032 .14 .002 .17 .002 .17 .10 H«O 6.9 3.61 7.2 1.60 7.2 4.17 7.8 5.5 7.8 2.61 CoO 11.4 .85 10.8 1.03 10.9 .51 11.9 .24 11.9 8.7 •OjO 2.20 1.84 2.79 1.64 2.72 1.47 2.18 .30 2.18 1.29 .18 2.17 .58 1.34 .58 2.59 .18 2.53 .18 .14 *2° 2 .19 .09 .21 .54 .22 .07 .14 .OO .14 .13 totol 99.0 82.5 98.4 74.7 97.8 81.3 99.3 78.6 99.3 81.5 00» 37 3328 45-1* 00» 37 Hl, 16-5 00» 37 334 17-1 00» 37' 335 10-4 Cl... »«l.fon siOj 51.0 36.3 51.0 43.3 51.3 42.7 50.1 41.6 TIOj 1.13 1.81 .13 1.22 .16 1.25 1.13 1.62 14.5 11.5 14.2 14.4 14.2 15.8 15.5 11.5 FoO1 10.3 15.3 9.9 14.5 9.9 15.0 9.4 14.6 (MO .17 .13 .16 .OO2 .20 .042 .15 .OO2 •to 7.8 3.42 8.3 3.72 8.2 3.04 8.5 3.28 CoO 11.9 7.4 12.7 .TI 12.6 1.11 11.9 .74 •OjO 2.18 .70 1.85 1.52 1.77 1.20 2.46 1.92 tjO .18 .23 .09 2.80 .08 2.27 .17 2.73 2 ••.0. .14 .05 .12 .06 .12 .07 .12 .OO tom 99.3 76.9 82.3 ».2 82.5 99.4 78.0 '»II ft rtportoo; •• F«O. 28tto« dtttetion Hatt. 5lro«n Mttriol *t |f«in Intorlor». Glo» »Mly»(« Ii OOP J7-..J2I 45-1 OOP 17-114 17-1 (km* - IDMT) i 1 • i.. I x J J TT nu nu i«n '•» the at» c— if at mt 10' OOP n-us 10-4 ODr* 37-3328 iS-1 •V - Mifflrl T in i T vi I y z z z I I I 1 T - I f I» KM 1M W Ml fUlf 4ll*f /•< «M #f» (•« »•!• Of Mgurt 62. Glatt neraal ixtd-palagenftt eo*pot c -» o o 133 1) In general, the element retention and depletion patterns are similar to the patterns seen in low-Fe, high-Si dredge sample palagonites; 2) In only one sample, ODP 37 332B 6-2, is there any true depletion of TiC>2 relative to parent glass (note that FeO is retained in this rind). This palagonite is also the most CaO depleted of any of the drill core samples. Diis correlation between TiO2 and CaO is not consistently seen in dredge samples; 3) AI2O3 is apparently depleted in the palagonite rinds of each of the samples relative to parent glass, this effect is greatest in ODP 37- 335 10-4 (the only ODP sample which has extensive zeolite formation) and the site 332B samples; 4) Like the dredge sample palagonites, MgO is more effectively retained than CaO, and K20 is highly enriched above the amcunt made available by glass dissolution; 5) Alkaline earth (MgO and CaO) concentrations are much higher in the palagonite of sample ODP 37 332B 45-1 (which is recrystallized to clay) than in other drill core sample palagonites or the low-Fe, high-Si dredge sample palagonites (except for MgO in USNM 110728- 468). Alkaline earth concentrations are also generally higher in the high-Fe, low-Si dredge sample palagonites. ODP 37 332B 45-1 palagonite is also not enriched in K2O, an oxide which shews strong enrichments in all other drill core samples. 134 Microprobe traverses of palaganite rinds were completed for samples ODP 37 332B 6-2, ODP 37 334 17-1, ODP 37 335 10-4 (zeolitized sample), and ODP 49 410A 2-2. Examples are given in Figures 63 through 68. The following can be concluded: 1) MgO decreases in concentration away from the glass/palagonite interface in all samples where traverses were run except ODP 37 335 10-4 (the sample with zeolites); 2) AI2O3 increases toward the outer portions of the rinds in sample ODP 37 332B 6-2 (the TIO2 depleted palagonite); 3) FeO increases toward the outer portions of palagonite rinds in samples ODP 37 335 10-4 and ODP 49 410A 2-2. The high iron portions of the rinds in ODP 49 410A 2-2 are distinctly redder than the inner portions of the same rinds; 4) Tetrahedral to octahedral cation ratios do not change in traverses across the palagonite rinds of all drill core samples. Values for this ratio remain consistently near 2.0 as expected for dioctahedral smectite (nontronite). The only deviation in this ratio occurs at the outermost (most FeO rich) portions of the rinds in ODP 49 410A 2-2 where the values dip progressively to approximately 1.7 and correlate well with increasing FeO content. Increasing Fe concentrations could be the result of excess Fe, perhaps in the form of an iron oxide or oxy-hydroxide phase, which is mixed in 135 ODP 37 332B 6-2 Traverse 1 24- 22 - 20 - * • • * , * • • * • • • 16- i 14 - 12 - i 0 10 i 1 • 8 • • * • 6 - 4 - 2 - 4 *444A4*4*. 4 4*4 0 20 40 60 80 PoaKion (micron») D AI2O3 • r«o • Moo 4 Coo ODP 37 332B 6-2 Traverse 1 4.S Poattlen (micron») • Fa • (V/VI Figur» 63. »atutt» of • nicroprob» trivartt of • paltgonitc rind in d«tp s»a drill core ••»pit OOP 37 332B 6-2. This is Travartt 1 of 6. Tht traverse begins in palagontte near the glast/palagonUe Interface and proceeds across the rind. The oxid* weight pereents shown in th* upper diagram »rt recalculated values, normal i ling to 100 X totals for the ten element-ox ides analyzed. The lower diagram gives the corresponding number of cations assuming a 22 oxygen (or 20 oxygtn, 4 hydroxyl) unit cell as well ss the resulting total tetrahedral/total octahedral cation ratios. 136 OOP 37 334 17-1 Traverse 1 24- 22 - a o o a a 0 20 - o ° a 0 • 0 18- a 4 » 4 4 4 4 * 18 • ' 4 4 4 14 • 4 a 12- 4 10 - 4 8 • 6- 4 • O t • 4 4 4 2- 4 4 é 4 C> 20 40 eo (micron*) 0 AJ2O3 4 W" • MgO A CaO OOP 37 334 17-1 Traverse 1 a o a a a : • t t • t • 20 40 eo Poaftlon (mlorone) ^ DM 4 r« « IV/Vl Figurt 64. Rttulti of a microprob* travtrtt of a palagonitt rind in dttp tta drill cor* tampI• ODP 37 334 17-1. This it Travtrta 1 of 6. The travtrt» bagint in palagonitt naar tht glatt/palagonit» intarfac* and procaada acrott the rind (which occurt in a fracture) to near the glatt/palagonit* interface on the other tide. Tht oxide weight ptrcentt shown in tht upper diagram arc recalculated valuta, normalizing to 100 X totala for the ten »I»merit-oxidet analyzed. The lower diagran givta tht corresponding number of cations assuming a 22 oxygen (or 20 oxygtn, 4 hydroxyl) unit ctll aa wall as tht resulting total tetrahedral/totat octthtdral cation ratios. 137 OOP 37 334 17-1 Traverse 6 •0 AJ2O3 ODP 37 334 J7—1 Traverse 6 PoaMlon (mierena) o Al • ?• • IV/Vt Figur* 65. Rt«utt» of a nfcroprebt tr«v*r*« of a palagonitt rind in dtcp tti drill core tamp 11 ODP 37 334 17-1. TM* it Trav«r»t 6 of 6. Tht travarst begins in pal««onit* n*jr th» gla»/palagonit* intcrfaea and proceeds across the rind (which occurs in a fracture) to near the glsss/palagonite interface on the other side. The oxidt weight percents shown in the upper diagram are recalculated values, normalizing to 100 X totals for the ten clement-ox i des analyzed. The lower diagram gives the corresponding number of cations assuming a 22 oxygen (or 20 oxygen, 4 hydroxyl) unit cell as well as the resulting total tetrahedral/total octahedral cation ratios. 138 ODP 37 335 10-4 Traverse 1 30 28 26 • • 24 •• ••••• 22 • • 20 no* BO a i = • "° o°0oQo0oo a°0 12 10 8 • • B « * • o o • 4 2 *V 0 20 40 60 90 100 120 140 160 PoaKion (mierena) O AI203 • r«0 • MgO A CoO ODP 37 335 10-4 Traverse 1 4.0 0.0 80 100 120 140 160 PeaKIon (rrWorone) A + to • IV/Vi Figura 66. Raaultt of a »icroproba travarta of a palagonita rind in daap laa drill core »ampla OOP 37 33S 10-4. This i» Travtraa 1 of 6. Tha travarta btgin-, in palagonfta nttr tha glatf/palagonita intarfaca and procaadt acrox tha rind. Tha oxida waight ptrcant» shown in tha uppar diagram tra recalculated values, normalizing to 100 X totals for tha tan elament-oxides analyzad. Tha lower diagram givas tha corrasponding number of cations assuming a 22 oxygen (or 20 oxygen, 4 hydroxyl) unit call as wall as tha resulting total tetrahedral/total octahedral cation ratios. 139 ODP 37 335 10-4 Traverse 4 x 28- 26- 24- 22- 20- 18 16 - 14 a o o a o a o ** a a o 12- 10 - e- 6- 4 • 2- 20 40 60 eo Poeltlon (mtcrona) O M2O3 CaO ODP 37 335 10-4 Traverse 4 4.0 IV/VI Ffgura 67. Rasultt of a »fcroprob» travtrst of a palagonita rind in datp *•» drill core •ampla ODP 37 33S 10-4. Thia ia Travarta 4 of 6. Tht travaria begins in palagonit* near tha glaaa/palagonite intarfaca and procetd» aero»» tht rind. The oxi da waight ptrcant» shown in tha uppar diagram art racalculattd valuas, normalizing to 100 % totals for tha tan altmtnt-oxides analyzed. The lower diagram gives the corresponding nuatotr of cations assuming a 22 oxygen (or 20 oxygen, 4 hydroxyl) unit cell as well as the resulting total tetrahtdral/total octahedral cation ratios. 140 ODP 49 41OA 2-2 Traverse 2 40- 38 - 30- a 0 K 25 - a o a o a a o o 0 a a a a 1 20 - o * o a a ° a a 15 - 0 i i • + * 10 - • • *> 5 - • • • # .••* • * • • * * *^? *• * V * ! i t 0 20 40 SO BO 100 120 P«*ttion (mlorone) a AI2O3 • r«o « M«O coo ODP 49 41 OA 2-2 Traverse 2 9.0 - 4.0 - a Q o a 0 S 3.0 0 a • 2.0 + • • • 1.0- • • • 0.0 20 BO 80 100 120 Poettl«n (miorone) O Al • re • IV/VI Figurt 68. ftttutts of a microprob» travers* of a pslagonite rind in deep saa drill core sample OOP 49 410A 2-2. This it Travers* 2 of 6. The traverse begins in palagonite near the glats/palagonite interface and proceeds across the rind (which occurs in a fracture) to n*»r the giass/palagonitc interface on the other side. The oxide weight pcrcents shown in the upper diagram are recalculated values, normalizing to 100 X totals for the ten element*ox i des analyzed. The lower diagrew gives the corresponding number of cations assuming a 22 oxygen (or 20 oxygen, 4 hydroxyl) unit cell as well as the resulting total tetrahedral/total octahedral cation ratios. 141 progressively higher proportions (relative to the nontronite component) toward the outer portions of these rinds. 3.3.2 Cementation Fine grained detrital calcite with abundant calcareous microfossils forms the matrix material in hyaloclasite samples ODP 37 332B 6-2, ODP 49 410A 2-1, and ODP 49-410A 2-2, completely filling all available space. In some areas in the matrix, the cement is recrystallized to larger calcite crystals. Clays occur in sample ODP 49-409 28-1 as fine grained, yellow-green colloforro masses between glass shards. A second clay cement is also seen, occurring as a red layer, 15 to 25 microns thick, with a well developed orientation perpendicular to substrate surfaces. This clay is the last cement formed, covering all exposed surfaces. Clay is abundant (not as a cement but as a replacement of palagonite) In sample ODP 37 332B 45-1. These clays are given in Table 17. The clay is chemically an Fe-saponite when excess Ca is eliminated from the calculated structural formula. Zeolite occurs as a minor cement in ODP 49-41QA 2-1, where radiating aggregates of phillipsite prisms are seen cementing one vesicle (Table 18 and Figure 69). Zeolites are more abundant as fracture fillings in ODP 37- 335 10-4. These zeolites are blocky and range from 10 to 50 microns in size. Microprobe data on these zeolites indicates that they are highly alkali rich with little calcium (Table 18 and Figure 70), probably phillipsite. 142 K ODP 49 41OA Z-1 Zeolites Na Ca+(Mg+Ba) Ffgurt 69. Plot of txeh§ngt«bl« citiont in philHp«U« in drill eort ••mple OOP 49 410A 2-1. 143 K ODP 37 335 10-4 Zeolites Na Ca+(Mg+Ba) ffgurt 70. Plot of txch«ng««blt cations in philltpsitt in drill eort tample OOP 37 33$ 10-4. 144 A minor amount of dark brown, opaque, Al-bearing manganese oxide is intergrown with the zeolite in fractures in ODP 37-335 10-4 and sane authigenic calcite can also be found in these fractures. The calcite is evidently the last cement to form. 3.4 AfflLYTT^If ELFTTFON MiCROSCOPy Analytical electron microscopy (AIM) has been used to examine selected dredge samples in order to establish the degree of chemical homogeneity and crystallinity of palagonite layers. Sample preparation (described previously) and analysis required an unusual amount of time. Thus, for the purpose of this report only two samples are disnisseri in detail. The analytical electron microscopy was completed by Dr. Takashi Murakami (on leave form the Japan Atomic Energy Research Institute). 3.4.1 SAMPIZ DESCRIPTION Sample USNM 113521-69 is a pillow rim glass from the Reykjanes Ridge (see Table 4). The glass is clear with low vesicularity and only a few microcrystals, plagioclase and olivine (Figure 71). External surfaces and fractures are coated by yellow to red palagonite which replaces the glass, forming rinds 10 to 100 microns thick (Figure 71). Rind thickness decreases as fractures penetrate the glass. Conpositionally the sample is a high iron (FEIT) glass (see Table 15}. The sanple is estimated to be .01 to 1.5 million years old based on its location relative to magnetic anomalies and age/depth 145 *- 100 pm Figure 71. Comparison of dredge samples chosen for AEM: a) thin-section photmicrograph of USNM 113521-69; b) thin section photomicrograph of USNM 113715; c) scanning electron photomicrograph of palagonite in USNM 113521-69, note the dissolution pitted glass surface (lower left); d) scanning electron photomicrograph of palagonite in USNM 113715. 146 relationships to the Reykjanes Ridge (see Iuyertiyk et al. 1978). Sample USNM 113715 is a pillow rim glass fran the Mid-Atlantic Ridge in the North Atlantic near the Kane Fracture Zone (Table 4). The glass has low vesicularity with scattered olivine microcrystals. The composition is similar to a typical Mid- Ocean Ridge Basalt (MCRB) (Table 15). Palagonite occurs as orange surface layers, replacing the glass on exterior surfaces and along fractures within the glass. The apparent rind thickness in thin section is 10 to 30 microns thick with clearly defined, multiple layers (Figure 71). The sample is less than 1 million years old based an its location near the center of magnetic anomaly 1 (Purdy, Rabinowitz, and Schouten 1978). Despite the similar compositions (Table 15) of both glasses and the similar appearance of the palagonite layers in thin section and. scanning electron micrographs (Figure 71), the compositions of the palagonites of each are distinctly different. Electron microprobe traverses across the layers reveal that SiC>2 in USNM 113521-69 is much depleted compared to that in USNM 113715 (Figure 60) and that FeO in USM* 113521-69 is much increased relative to AI2O3 compared to that in USNM 113715 (Figures 54 through 59). Also, in traverses across the layers, the electron microprobe analyses provide only the slightest hint of compositional zoning and heterogeneity within the layers (Figures 54 through 59). 147 3.4.2 ANALYTICAL EEECIPON MKROSODP* 3.4.2.1 OSNM 113521-69 Three zones of distinctly different texture, erystallinity and caqposition are apparent in this sample (Figure 72a). Transmission electron micrographs and electron diffraction patterns of each of the zones are shown in Figures 72 through 74. Three phases are identified, and they are labelled according to the order of their apparent paragenesis (e.g., A, B, and C). In Zone 1 (nearest the glass/palagonite interface), the dominant phase is Hiase c consisting of mats of thin fibers. Fhase C is accompanied by poorly crystalline phases (A and B, described below) which are microstucturally similar to what is observed in Zone 2 (Figures 73b and 73C). The fibers (phase C) are two to ten nm in thickness (approximately two to ten unit layers), and less than 50 nm in length (Figures 72c and 72d). They are braided (Figure 72c) and, in some areas, form an interlocking network (Figure 72b). Diffuse diffraction rings indicate that these fibers are not well crystallized. There are two d-spacings for the lattice fringes of the fibers: 0.6 to 0.7 nm (black arrow in Figure 72d) and 1.2 to 1.4 nm (white arrow in Figure 72d). More accurate d-spacings could not be obtained because the much brighter incident beam obscured the intensities of inner spots and rings. Zone 2 (Figures 73b, 73c, and 73d) is the least crystallized area of the layer. However, diffuse rings (Figures 73b and 73c) and the dark field image (Figure 73d; Figure 73c is a corresponding bright field image) show the crystallinity in Zone 2. A poorly crystalline phase (A) with a mottled 148 Figure 72. AEM photomicrographs of deep sea dredge sample USNM 113521-69: a) low magnification view of entire palagonite layer with indicated zones (Zone 1 is nearest glass); b) high magnification view of fiber (phase C) network in Zone 1; c) braided fibers in Zone 1; d) high resolution photomicrograph of Zone 1 fibers. Black arrow indicates 0.6 to 0.7 nm lattice fringes and the white arrow indicates 1.2 to 1.4 nm lattice fringes. 149 Figure 73. AEM photomicrographs of sample USNM 113521-69 palagonite Zone 2: a) transition zone between Zone 1 and Zone 2; b) Phase A showing mottled diffraction contrast and poorly crystalline character; c) interlocking network of crystals in Phase A; d) dark field image corresponding to image c. 150 Figure 74. AEM photomicrographs of sample USNM 113521-69 palagonite Zone 3 (farthest from glass): a) transition zone between Zone 2 and Zone 3; b) fibers with 0.6 to 0.7 nm lattice fringes (black arrow) and 1.2 to 1.4 nm lattice fringes (white arrow); c) fibers with 0.6 to 0.7 nm lattice fringes; d) high resolution image of fibers in image c. 151 diffraction contrast (Figure 73b) forms an interlocking network of crystals developing in the matrix of phase A (Figure 73c). The fibers are less abundant in Zone 2 than in Zones 1 and 3. There is a transition zone (approximately 1 micron) between Zones 1 and 2, in which all the microstructures observed in Zones 1 and 2 are present (Figure 73a). Zone 3 (Figures 74b, 74c, and 74d) is sLnilar to Zone 1 with respect to its microstructures. Fibers with 0.6 to 0.7 ran lattice fringes (black arrow in Figure 74b) and 1.2 to 1.4 nm lattice fringes (white arrow in Figure 74b) dominate the microstructure. The distinct features in Zone 3 are the abundant fibers v\th 0.6 to 0.7 nm lattice fringes (Figure 74c). The fibers are well developed, up to 10 nm in thickness and 50 nm in length (Figures 74c and 74d). A transtiion zone exists between zones 2 and 3 (Figure 74a). This transition zone, like the transition zone between zones 1 and 2, contains all the microstructures observed in Zones 2 and 3. Approximately 60 points were analyzed by EDS in phases A, B, and C. Each phase shows compositional variations. Table 20 gives the average compositions of phases A, B, and C. Note, that due to the small size of the network- forming phase (C), the analyzed areas include a mixture of phases B and C. Compositional differences may be characterized as follows: 1) Fe and si are higher, and Al is lower in phase C than phases A and B (FeO, SiO2, and AI2O3 concentrations (wt.%) are 48, 22, and 13 in phase C; 41, 17, and 22 in phase A; 46, 14, and 21 in phase B, respectively) 152 Tablt 20. Compositions (wt.X) of phastt A, B, and C of USNM 113521-69 datarainsd by AEH. Phat* A Phas* I Phast C Zont 1 Zona 2 Zon« 3 17(4)" 14(3) 22(5) 25(4) 19(4) 19(4) sfo2 TiO2 5(4) 4.9(8) 6(1) Al2°3 22(5) 21(5) 13(3) 13(2) 13(5) 11(2) FtO2 41(4) 46(5) 48(7) 44(4) 49(6) 54(6) MgO 2.7(9) 2.7(7) 1.4(8) CaO 0.6(3) 0.4(3) 0.3(3) K2O 0.1(1) 0.03(6) 0.1(1) P2O5 4(1) 5C) 2(1) S0 2.0(5) 1.6(8) 0.9(8) 3 3 4(3) 4(J) 8(3) Concantratfont of SiQ^, *l2°3' *nd f'° 'ton* *r* fliv»n to show variation aMong tht lonts. 2All iron rtportsd as F«O. Contamination of Cl fron tht tpoxy rtsin. Ona standard deviation for last digit is givtn in parenthtsas. 153 2) The composition of phase C varies among the zones; FeO concentration increases (44 to 54 wt.%) and SiO2 and AI2O3 concentrations decrease (25 to 19 wt.% and 13 to 11 wt.% respectively) in Zone 3 (nearest the solution). Because phases A and B both have d-spacings of approximately 0.26, 0.21 and 0.18 ran and phase B forms in phase A, phases A and B are probably the same. In addition, the similarity in the electron diffration patterns (powder rings of phase C: 0.65-0.70, 0.34, 0.25, 0.20, 0.18, and 0.16 ran for d- spacings; diffuse rings of a mixed phase of A and B: approximately 0.65m 0.43, 0.33, 0.27, 0.20, 0.18, and 0.15 ran for d-spacings) and in compositions (Table 20) strongly suggest that phase C is also the same as phases A and B. Figure 75 shows a typical microstructure in Zone 2, in which phase C is present only in the channel running parallel to the zones, and phase C extends into the channel. This texture suggests a sequence of phase development of first A, the B, and finally C. The channels act as paths for cation transport. Based on the electron diffraction data and the composition, these phases are smectites, other minor crystalline phases included monönorillonite, goethite, gibbsite and serpentine. These phases are usually found in Zone 1. 3.4.2.2 USNM 113715 The layer formed in this sample is very different to that observed in the previous sample, despite the similarity in alteration environment and bulk glass composition (Table 15). The entire layer is shown in Figure 76a. The 154 figure 75. AEH photomicrograph of Zone 2 in USNM 113521-69 which contains both Phase B and Phase C. Phase C is seen only in the channel (upper left to lower right) which runs parallel to the zones (parallel to the glass/palagonite interface). 155 Figure 76. AEM photomicrographs of palagonite in deep sea dredge sample USNM 113715: a) low magnification image of the entire layer; b) high magnification image of the area of the rind nearest the solution (farthest from the glass), note that there are no diffraction maxima in the electron diffraction pattern indicating that the palagonite is amorphous in this location; c) fibers in an amorphous matrix in the middle portion of the rind; d) the amorphous inner zone (nearest to the glass). 156 layer is essentially amorphous (Figures 76b and 76d). There is no indication of crystalline phases in the electron diffraction patterns (Figures 76b and 76d) for the innermost area (one micron in thickness) and the outermost area (two microns in thickness). In the rest of the layer, the poorly developed fibers in an amorphous matrix give only a diffuse diffraction halo (Figure 76c). The innermost and outermost amorphous areas and the interior fiber-network areas have different compositions (nos. l and 3, and no. 2 in Table 21, respectively); the interior areas are poorer in Fe and richer in Si than the amorphous areas (FeO concentration is 15 wt.% in the latter, and approximately 30 wt.% in the former; SiO2 concentration is 50 wt.% in the latter, and approximately 20 wt. % in the former). 3.4.3 CONCLUSIONS There are no generalizations that may be made from this preliminary work; however, there are two important observations. First, palagonite layers that appeared to be amorphous and chemically homogeneous when examined in thin section or by scanning electron microscopy and electron microprobe analysis display a distinct zonation in terms of the degree of crystallinity, the microstructures, and the chemical composition when examined in finer detail by analytical electron microscopy. Second, samples of similar age and from similar alteration environments show dramatic and easily detectable differences in their zonation when examined by MM. Observations of experimentally altered basaltic glasses (Crcvisier et al. 1983; Crovisier, Honnorez, and Qerhart 1987) by AEM indicate the formation of a layered structure in the alteration rind. This structure is the result of 157 Tablt 21. Compositions (wt.X) of difftrcnt areas of USNM 113715 determined by AEM. Arta 11 Arta 22 Area33 6 sio2 25.8 50(7) 17.5 TiO2 5.1 2(1) 4.6 Al2°3 29.6 29(3) 26.4 FeO4 29.3 15(6) 33.0 MgO 2.3 1.3(6) 1.6 CaO 1.3 0.3(3) 0.5 K2O 0.1 0.2(4) 0.0 P2O5 2.8 0.7(5) 3.0 S03 . 1.4 0.2(2) 3.0 ci,o«5 2.4 KD 11.1 11nnermost area nearest to glas* (Figure 76d). 2Arti except innermost and outermost areas (Figure 76c). 'outermost area nearest to solution (Figure 76b). All iron reported as FeO. 'Contamination of Cl from the epoxy resin. °0ne standard deviation for last digit is given in parentheses. 158 mineralogical differences which include (from the outermost surface of the rind toward glass) saponite, hydrotalcite and a high Fe and Mg clay mineral. The detailed analysis of leached surfaces by analytical electron microscopy should be a standard procedure in their characterization and in phase identification. 4 DISCUSSION 4.1 PAIAGONTTE Dibble and Tiller (1981) have modeled the formation of alteration products as following an Ostwald step sequence, a series of irreversible reactions which sequentially form more thermodynamically stable phases. They advocate kinetic controls and state: "... metastable reactions occur because formation of less stable phases can lower the total free energy of the system faster than growth of the stable phase assemblage." At high dissolution rates, the near-glass environment may became supersaturated with respect to smectites or aluminnsilicate gels. These fast growing phases contain larger masses of Si and Al per unit growth time, compared to thermodynamically more stable, isochemical silicates (Dibble and Tiller 1981). The formation of gels, such as palagonite, can be envisaged as rapidly precipitated crystallites (10 to 100 angstroms in radius) which represent critical nuclei with essentially no crystal growth under conditions 159 of high supersaturation (see Berner 1971). If the dissolved aluminum species is predominantly Al^, the precipitate should be clay crystallites (DeJong, Schrann, and Parziale 1983). Nontronite is the expected product at low temperatures under oxidizing conditions (Kurnusov 1982) and, in fact, palagonite (except for that in high-Fe deep sea dredge samples) appears to be a dioctahedral, Fe-rich clay-lite material (see, for example, Figures 17 through 23). DeJong, Schramm, and Parziale indicate that Al71 should be the predominant form of aluminum in acid to slightly alkaline solutions, up to a pH of approximately 8.0. Al^ (as A1(OH)4~) should be the only aluminum species present in solution at high pH (or high hydrolysis number, that is, CH/A1 from 5 to 15). The formation of palagonite in highly aDcaline conditions would require the presence of sLrony electron acceptors (e.g., Mg2"1") which could cause reversion of the aluminate tetrahedral cluster to Al71 (DeJong, Schramm, and Parziale 1983). A possible indicator of palagonite formed under high pH conditions (above 8.0) might, therefore, be low total aluminum content in general, and low Al^1 in particular, with octahedral sites dominated by Fe and Mg. Ihis assumes that the aluminum which is released as the glass dissolves does not retain a metastable configuration resulting from partial polymerization of the glass progenitor. In alkaline seawater, the Si content of palagonite should also be relatively high compared to aluminum, similar to the Si/Al values in deep sea phJlipsite as compared to non-marine phillipsite (Fisher and Schmincke 1984). Figure 77 plots the numbers of si and Al^ cations in palaganites, calculated assuming a clay mineral (22 oxygen) structure. Seawater altered glasses have formed palagonite which is generally lower in Al^ and higher in Si than subglacially produced glasses palaganitized in fresh water. No+i» that deep sea drill core palagonites do not reach the very low aluminum values seen in some dredge sanples or sane 160 Cation Proportions iri Palagonite 2.50 q ° - Iceland (sw - aeowoter altered) D m Dredge samples o * • Drill tor» SamplM to 2.00 v X o 1.50 i es VI .2 1.00 "o5 > 0.50 ; seawater altered 0.00 0 4.50 5.00 5.50 6.Ö0 6.50 7.00 7.50 Si (cat!ons/22 oxygens) Figurt 77. Octahedral aluminum (AlVI) va. Si cation content of palagonite calculated aaauning 22 oxygen*. An »tu» above a data point indicate* that it represents a seawater altered sample from !:«land. Filled symbols indicate zcolitized samples. 161 seawater-altered samples from Iceland. The lowest Al^ palagonites from Iceland are the seawater altered samples from Kopavogur and Reykjanes. Figure 78 shows the average octahedral cation proportions of palagonites assuming a clay mineral structure. Nearly all palagonites analyzed are dominantly Fe in octahedral sites. The exceptions being three fresh-water altered Iceland palagonites which are aluminous (M-l, M-7, and R-l) and one deep sea dredge sample palagonite (USNM 110728-468) which is dcciinantly Mg in octahedral sites. Fresh water formed palagonites contain at least 20% Al in octahedral sites. Seawater formed palagonites do not exceed about 40% Al and go as low as 0% Al in octahedral sites. Deep sea drill core samples tend toward higher magnesium and lower iron when compared to other seawater altered samples. According to solubility-pH relations defined by Lougnan (1969), Reesman, Pickett, and Keller (1969), and Baes and Meaner (1979), the point of minimum solubility (isoelectric point) for aluminum occurs at a pH of about 6.7. The total concentration of aluminum in solution should increase toward both higher and lower pH values (away from the isoelectric point). Thus, those palagonites which have retained alumina (most of the Pleistocene subglacial samples from Iceland) may have formed in near-neutral solutions. Depletion of aluminum in palagonite relative to glass could result from palagonitization at either low pH or high pH. The pH of lake waters in volcanic craters has been measured in at least one location to be about 1.5 (Rankama and Sahara 1950). Increased acidity in groundwater prior to eruptions of Kraf la in Iceland have also been reported (Bjornsson et al. 1977). Bxjlacial lake waters (lakes created by subglacial volcanism) could, therefore, have relatively low pH 162 Mg+(Mn) Palagonite Iceland Drtdge Samples octahedral cations Drill Cora Samples Fe+(Ti) Figurt 78. Octahedrat cation plot of patagonites assuming 22 oxygens. An "sw" above a data point indicates that it represents a >eawater altered sample from Iceland. The shaded area includes all frefh water altered samples. 163 values, particularly early in their development. Low aluminum palagonites from subglacial volcanics oculd have formed in these low pH conditions. Falagonite has been formed experimentally at a pH of 6 in fresh water at 25°C (Fumes 1975). The duration of englacial lakes is not known but could not have been longer than the duration of the ice sheets in which they formed. Subglacial deposits in Iceland have been exposed to meteoric water for at least the last 8000 years. Solutions affecting deep sea samples may be either alkaline or acid. Interstitial water studies indicate that pore waters in deep sea sediments can be at least as alkaline (pH from 7.82 to 8.72, Miller, Lawrence, and Gieskes 1979) as bulk seawater (pH=7.8). Hydrothermal vent water at deep sea sites in the Guaymas Basin (Gulf of California) have reported end member pH values from 3.3 to 3.8 (Von TVnrcn et al. 1985) and sediment pore waters frorn this area have a pH as low as about 6.0 at DSDP site 477 ranging up to 8.15 at DSDP site 478 (Gieskes et al. 1982). Bischoff and Dickscn (1975) envision the following scenario for hydrothennal systems in the oceanic crust based on experiments: 1) Initial uptake of Mg and CH from seawater to form montmorillonite lowers the Mg content cf the solution and lowers the pH from an initial value of 8.1 to 4.5 at 200°C (the measured pH at 25°C is about 4.0); 2) As Mg is used up, the pH rises to above 5.0, and Fe (as Fe2+), Mn, and Si are released into solution; 3) As the solution rises and cools, Si is precipitated, Fe and Mn solution concentrations do not change appreciably; 164 4) When the cool, reducing solution encounters the oxygenated bottom waters of the open ocean, Fe2+ is oxidized to Fe3+ and is precipitated as colloidal hydroxy-hydrates. Clearly, sub-seabed pore waters can vary considerably in pH, being easily modified by such factors as hydrothermal alteration of minerals as well as degradation of organic matter. Additionally, glasses at the ocean floor may be exposed to significantly modified, metal-rich seawater while still in the vicinity of the mid-ocean ridge. There is no simple environmental constraint which prohibits either acid or alkaline conditions from producing the low aluminum palagonites in deep sea drill core samples. Dredge samples are less likely to have been exposed to low pH water as they have existed an the surface of the sea floor, exposed to the open ocean, throughout -cheir history. Aluminum depletions in low-Fe, high-Si palagonites, therefore, are likely due to glass alteration under alkaline conditions. The dissolved Vq concentrations in interstitial waters of deep sea sediment tend to decrease with depth below the sediment/water interface with a corresponding increase in concentration of Ca (Miller, Lawrence, and Gieskes 1979; Gieskes and Reese 1979; Gieskes et al. 1982). Gieskes and Peese (1979) propose that the source of these gradients is basalt weathering which results in net uptake of Mg (as sapanite) and release of Ca. Lang-term alteration of dredge samples could, therefore, produce palagonite lower in Ca than drill core samples altered for long time periods beneath the sediment cover. This is because the Ca release rate should be higher in an open ocean environment where there is no sediment cover to inhibit diffusion. In fact, there appears to be no consistent difference in the Ca contents of dredge and drill core 165 sample palagonites. Mast of the palagonitizatian of drill core samples may have taken place while those erupted materials were either still exposed to the open ocean or when sediment covers were not yet thick enough to affect interstitial water compositions. One of the major differences between fresh water formed and seawater formed palagonites is the higher concentration of alkalis, particularly K, in seawater palagonite. Potassium is enriched in deep sea palagonites to values as much as nearly 200 times that originally present in the glass. Sodium is either conserved or slightly depleted in deep sea palagonites relative to parent glass. Fresh water alteration tends to produce palagonites which have alkalis present in concentrations below one tenth the amount in fresh glass. Drever (1982) indicates that average river water contains 5.15 mg/kg Na+ and 1.3 mg/kg K*", and that average seawater contains 10,760 mg/kg Na+ and 399 mg/kg K4. If palagonite composition reflects solution composition, then seawater formed palagonite should be higher in total alkalis than fresh water palagonite. Hie alkali content is due to interlayer occupancy in the clay structure and, in the case of colloidal material such as palagonite, also to physical adsorption onto particle surfaces. Eberl (1980) suggests that the lower exchange free energy of X4" relative to Na+ causes a slight preference for the K4" cation in clays. High layer charge, however, can cause a rapid increase in K*" selectivity over Na+ due to the lower dehydration energy of K4". Palagonite never exhibits large iron losses, indicating that most iron is probably in the less soluble ferric state. This is consistent with the interpretation of palagonite consisting primarily of nontranite crystallites, based on cation proportions. 166 is generally retained or even enriched during palagonitization, in accordance with the insolubility of the non-hydrated ion. In some deep sea samples (the high-Fe, low-Si palagonites), TiO2 is enriched above the amount made available by glass dissolution by about the same amount as FeO. Apparent titanium enrichment could result from passive accumulation only if the volume of glass reacted was actually greater than the volume of palagonite (that is, the reaction was not iscvolumetric). Petrographic evidence suggests that there is no shrinkage during palagonitization with the possible exception of the high-Fe, low-Si deep sea dredge palagonites. Sample M-l from Iceland, low-Fe, high-Si dredge samples USNM 113715, USNM 114045-3A, USNM 111238-1, and deep sea drill core sample ODP 37 332B 6-2 all have palagonite rir.'i that are depleted in TIO2 relative to glass. However, iron, in these samples was completely reta .ined during palagonitization, even showing enrichment in sane cases. If titanium were truly insoluble and was conserved during palagonitization of these samples, then the rinds would have to be of higher volume or density than the reacted glass. Volume increase would be readily discemable in thin section and is not seen in any of these samples. Analytical totals for these palagonites are from 70 to 88 wt. %, the deviation from 100 % being do to the presence of water. The palagonite density must, therefore, be considerably below the density of glass, as suggested free density measurements of Hawaiian palagonites (densities from 2.75 to 2.85 g/cm3 for basaltic glass, and from 1.93 to 2.14 g/cm3 for palagonite [Hay and lijima 1968a; 1968b]). The high water content of palagonite has been established by thermogravimetric analysis (Allen et al. 1981). 167 If shrinkage of palagonite relative to the volume of reacted glass were responsible for the apparent titanium and iron enrichments in the high-Fe, low-Si deep sea dredge sample palagonites, then the Si losses would be 2 to 5 tines greater than calculated for isovolumetric reaction. If Si losses were similar to other deep sea dredge samples then the observed Fe and Ti enrichments are real. Note that some Ti is actually lost fran some of the low-Fe, high-Si deep sea dredge palagonites indicating that Ti goes into solution. Titanium mobility may be related to its structural environment within the parent glass. Titanium in rutile or ilmenite structures will not be released during hydrolysis of the glass as Ti-oxides are relatively stable (Rankama and Sahara 1950). In such a case, Ti may remain in the insoluble residue as Tio2. If, however, the titanium is in pyroxene or olivine structures, then it may be released and hydrolyzed as Ti(0H)3 when the relatively unstable ferromagnesian-silicate structure is broken down. The titanium depleted palagonites are generally restricted to deep sea dredge sanples. There is no reason why dredge sample glass should differ structurally from drill core sample glass. Glass composition, depth of formation beneath the sea, and eruption temperature will be approximately equivalent for each of these glasses. The behavior of titanium is, therefore, most likely due to properties of the aqueous environment. In some cases, palagonite is compositionally heterogeneous. The major element gradients detected by microprobe traverses of rinds indicate that this heterogeneity is systematic. Some possible explanations for these gradients include: 1) thft increasing difficulty in diffusing dissolved material through 168 a thickening rind during palagonitization; 2) sane response to changes in post-palagonitization chemical potentials or 3) environmental changes during palagonitization, affecting solubilities of species and, consequently, product composition. If gradients in palagonite were purely the result of inhibited diffusion, then the gradients would always be present, always be in the same direction, and would be quantitatively related to rind thickness (the effect should be greater for thicker rinds). None of these conditions are met among the gradients detected in the palagonite rinds in samples from Iceland or the deep sea. Gradients clearly established in rinds in one sample may not exist in other samples. The trends for AI2O3 and FeO are in the opposite directions in Reyk.-l compared to the general directions noted in Kb.-l or A-l. Cation proportions indicate that the tetrahedral (TV) to octahedral (VI) ratio is more or less constant across palagonite rinds, even where strong compositional gradients exist. Post-palagonitization effects should not alter the tetrahedral or octahedral cation proportions without reconstructive transformation. Iron and aluminum are not exchangeable elements in a silicate structure such as clay. After crystallization, the only mechanism for changing aluminum or iron concentrations in the clay is recrystallization. Only the exchangeable cations, Na, K, and Ca, can be added or removed from palagonite in order to satisfy exchange equilibria. Compositional gradients in x-ray amorphous, non-birefringent palagonite must be at least partially the result of the palagonitization environment. Solubility relationships will change with time as glass hydrolysis proceeds 169 because solution pH increases and dissolved element concentrations change as they are released or ccnsunwri. Die greater the reck/water ratio, the greater the magnitude of s>\/i effects. If flow rates through the deposit are more rapid than eler • * il release rates (low effective surface area/solution volume) thfj- solution oonoentrations may not change significantly. Gradients are seen in palagonite rinds from Tungufell, Iceland. The ,4dients exist in rinds associated with both zeolite-filled and clay-filled vesicles. Outermost (first-formed) palagonite, which is furthest from the glass/palagonite interface, is higher in Fe and Mg, and correspondingly lower in Al and Ca, than inner portions of the rinds (see Figures 17 through 23). The concentrations of these elements change systematically toward the inner (later-formed) part of the rind. In terms of cation proportions, Fe is more abundant than Al in the outer portions of the rinds. Traversing toward the glass/palagordte interface, the cation proportions change, with Al eventually becoming more abundant than Fe. This changeover occurs at a riisfannp of 20 to 40 microns from the interface. In rinds associated with zeolite filled vesicles (rinds 70 to 190 microns thick), MgO is higher than CaO in outer portions of the rinds with CaO eventually becoming more abundant moving toward the glass/palagonite interface. In clay-filled vesicles (rinds 25 to 40 microns thick), the same trends are seen with MgO and CaO except that CaO is always more abundant. For all elements, the gradients in rinds associated with clay filled vesicles are nearly identical in slope and overall concentrations to the Innermost 25 to 40 microns of the gradients associated with zeolite filled vesicles (see Figure 23). In some of the rinds associated with zeolite filled vesicles, sharp decreases in Fe and Ti, with a corresponding increase in Mg, occur in the outermost 20 to 30 microns. 170 If palagonitization in the two vesicle types started with the sane solution concentrations, then the gradients would be expected to match over their outermost (first-formed) portions. That they, in fact, match over their innermost (latest-formed) portions is perplexing. The conditions at the time palagonitization began in the day-filled vesicles apparently matches the conditions near the end of palagonitization of the zeolite-filled vesicles. Either solution compositions in the two vesicle types were initially different or the gradients are somehow the result of syn- or post-palagononitization effects such as cementation authigenesis. Mass balance calculations were performed in order to evaluate if there had been significant mass loss from palagonitized vesicles or mass input into vesicles. These calculations asamp spherical vesicles and are based on the observations that palagonitization was isovolumetric in Tungufell samples, and that zeolite or clay tends to completely cement the vesicles in which they occur, leaving no pore space (see Figure 15). Figure 79 shows the difficulty of making volume estimates for vesicles based on thin section measurements. The relative volume of palagonite will always be somewhat overestimated relative to zeolite or clay cement. By performing the calculations on many vesicles, some consistancy is attained. Figure 80 shows the results of such mass balance calculations for zeolite-filled vesicles. Si and Al are retained, or are perhaps somewhat enriched, relative to the amount made available by glass dissolution. Ti and Fe are enriched, and K is highly enriched, indicating some mass input. There is no compelling evidence for shrinkage during palagonitization (glass and palagonite remain in contact independent of rind thickness; at intersections of fractures with vesicles, 171 PALAGONITE RIND true diameter measured measured diameter thickness THIN SECTION true CUT ~ thickness VESICLE GLASS Figure 79. Illustration of tht difficulty in obtaining volume information from vtsicles in thin ••ction. Th« measured palagonita thicknttt Mill bt grtattr than the trut thickntss unless tht cut of tht thin atetion happt a to coincide with the largtst dimension of th* associated vtaiclt. Similarly, tht vtticlt diameter measured Mill bt somewhat Itta than tht tru- diameter. If tilt trut vtticlt sizes tend to be similar, then rind measurements on vtaiclts which art appartntly larger will bt eloatr to trut valutt. 172 T-2 Vesicle Masa Balance (with zeolites) -100 Si Al Ft Mn Mg Ca Na .iOK P Figure 80. Calcultttd »••» balanct of ztotitt-filltd vtsiclt* from Tungufcll, Icttand sample T-2. Tht valu* dtlta it tht molar conetntration of an altmtnt in the alteration products minus the amount available from glass dissolution, such that negative values indicate net system loss. The length of the bars represents the 1-sigma standard deviations. Elements not represented with bars are elements not present in zeolite. Note that K values are shown in units of .10K. 173 the fractures do not hporaip open as they are palagonitized). Even assuming that the Ti enrichment is due to shrinking and revising all masses downward accordingly, K is so highly enriched that mass input is still required. Figure 81 shows the mass balance results of day-filled vesicles compared to zeolite-filled vesicles. Clearly, mass input is required to form the clays in these vesicles. Such a high degree of mass input is not indicated for the zeolite-filled vesicles, which perhaps require only K-rich initial solutions to produce all alteration products. High degrees of mass input, as seen in the clay-filled vesicles in Tungufell samples, indicate that flow rates must have been high in these vesicles compared to the zeolite-filled vesicles. Note that the K content of palagonite associated with clay-filled vesicles is higher than the K content of palagonite associated with zeolitized vesicles. Externally buffered potassium concentrations (open system) could allow higher K concentrations in alteration products of clay-filled vesicles than the more closed system scenario (zeolitized vesicles) because the zeolite-filled vesicles would have had only the initial amount of potassium available. Note that there is net system magnesium increase (Figure 81), primarily due to clay precipitation. However, the MgO content of palagonite is actually lower in clay-filled vesicles compared to zeolite filled vesicles. This may indicate that clay formation took place contemporaries isly with palagonitization in the clay- filled vesicles. This clay precipitation, with net Mg deposition, may have locally affected pH, causing less alkaline conditions in these vesicles (net deprotcnation of the initially hydrolyzed Mg cation). This would explain why zeolitized vesicles have thicker palagonite rinds as glass dissolution and zeolite formation is promoted at higher pH. 174 T-2 Vesicle Mass Balance 400: : C c C* • Z a zeollie filled C * clay filled 300 C c "c i | • c O 200 m a < c z "1 O1 °° ; z z_. i 1 z c "" • z_ z "I- Si i4( Fe Mn Mg Ca Na .101 P Figur* 81. Calculated mass balanc* of ztolitt and clay-filltd vtsiclts from Tungufell, Icaland sa*pl« T-2. Th« valu» dalta is th» molar concantration of an element in th* altaration products minus tha amount available from glass dissolution, such that negativa values indicate net system loss. The length of the bars represents the 1-sigwa standard deviations. Elements not represented with bars are elements not present in zeolite. Note that K values are shown in units of .10K. 175 The remaining question for the Tungufell samples is why some vesicles experienced more open system conditions (higher flow rates and mass input) than others. The answer may lie in the fracture systems which initially allowed solutions tc reach the vesicles. Thin sections reveal that the fractures which reach zeolitized vesicles are few in number and appear closed compared to fractures entering day filled vesicles (clay cements these fractures as well). There are no zeolitized fractures. This is consistent with the data presented in Figure 15. Under open system conditions, there is no relationship between rind thickness and vesicle size. Under restricted flow conditions (zeolitized vesicles), palagonite rind thickness increases with apparent vesicle size. The thickness of the palagonite rind under closed system conditions in a vesicle will be dependent upon the length of time required to reach surface silica saturation. Smaller vesicles will require less time. Surface Si-saturation may not be maintained or possibly even reached as zeolitization results in net consumption of Si. The increasing Al and Ca concentrations relative to Fe and Mg concentrations of the Tungufell palagonite rinds (regardless of cement type) may be explained by solution composition changes which occur during palagcnitization in response to palagcnitization and authigenesis. If the gradients are due to the growth of aliiminosilicate authigenic cements, which might deplete the outer portions of the rinds in Al and Ca (in the case of zeolite growth), then the gradients should somehow reflect the amounts of elements required to grow the cement. For clay-filled vesicles, even more Al appears to be required to crystallize the required volumes of cement than is required to form zeolite in zeolitized vesicles (Figure 81). If Al was being 176 extracted from the rinds, then those associated with clay-filled vesicles would be more depleted than those of zeolite-filled vesicles. The resulting Al gradients across the day-associated rinds would resemble the outer portions of the rinds associated with zeolite-filled vesicles. As noted above, the gradients in rinds associated with day-filled vesicles most resemble the gradients of the inner portions of the rinds in zeolitized vesides. The most plausible explanation for the presence of chemical gradients is that they represent changing solution compositions with time. Compositional gradients are also seen in seawater altered sanqples from Iceland. Despite their similar ages and alteration environment, thes* samples differ remarkably from each other in terns of rind thickness, rind composition, and cement type. The deposits at Reykjanes and Brimnes are pillow basalts with interpillow hyaloclastite and the deposits at Kopavogur and Arnarnesvogur are hyaloclastite breccias. Kopavogur, Arnarnesvogur, and Reykjanes palagonites are all depleted in aluminum relative to parent glass. Palagonite composition varies systematically across rinds in samples from Kopavogur and Arnarnesvogur though sane trends can be seen in samples from Brimnes and Reykjanes as well. All palagonite rinds from seawater altered samples from Iceland show a tendency for high concentrations of Mg in the outermost (first-formed) portions of palagonite rinds relative to portions of the rinds closer to the glass/palagonite interface. This effect is strongest in samples from Kopavogur, which also show a strong corresponding increase in Ca toward the interior (later-fornwd) portions of the rinds. There is a suggestion of higher Al and lower Fe toward the outer portions of rinds from Reykjanes. This is opposite to the trends seen in palagonite rinds from Tungufell and in rinds from Kopavogur and Arnarnesvogur. Aluminum increasing 177 with di stamp toward the glass/palagonite interface is only weakly developed in Kopavogur palagonite rinds. No such systematic change in Al and Fe can be discerned in samples from Brimnes. Zeolite accounts for approximately 10 volume percent of the Köpavogur deposits (pore space 25 volume percent) and .40 volume percent of Arnarnesvogur deposits (pore space 30 volume percent). In both cases, zeolites overgrow earlier formed day's. Zeolites are absent in samples from Reykjanes and only some rare phillipsite crystals are seen in Brimnes samples. With reference to the Tungufell samples discussed above, the presence of zeolites may be indicative of closed system conditions. Thus the gradients in Kopavogur and Arnarnesvogur palagonite rinds may result fron changing solution compositions. The solution composition will evolve as glass hydrolysis proceeds under closed system conditions. Open system alteration at Brimnes allowed the deposition of abundant Fe-rich clay and did not produce aluminum depletion in the rind. Many of the differences in the seawater altered samples from Iceland could be due to the effects of groundwater outflow and mixing of seawater and fresh water. Lower salinity is correlated with higher dissolved aluminum concentrations in near-shore marine environments such as the East China Sea (Mackin and Aller 1984) and the Tamar Estuary of southwest England (Morris, Howland, and Bale 1986). Similar effects in the restricted marine inlets at Kopavogur and Arnarnesvogur may have caused significant changes in the composition of seawater which altered the deposits at these locations. 178 Compositional gradients in the low-Fe, high-Si deep sea dredge sample palagonite USNM 113715 are iiiconsistent and no general trends are evident. This is likely due to the open system conditions of the ocean to which most dredge samples have been exposed throughout their history. Gradients in the high-Fe, low-Si dredge sample palagonite USNM 113521-69 indicate a high-Fe (and corresponding low-Al) region in the outermost palagonite. AEM evaluation of this sample indicates both silicate (clay- like) and Fe-oxide components in this rind. Thus the higher iron areas detected by EMPA may represent higher relative proportions of Fe-cxide to smectite clay. The cause of variation in the Fe-oxide proportions of the rind may be related to the environment which forms such generally Fe-rich palagonites in the deep sea. As previously mentioned, correcting for shrinkage (or lower density) to allow titanium conservation with no enrichment in these samples leads to the conclusion that Si losses would have been extreme, perhaps to one-tenth of the amount originally present in the glass. The only other interpretation of these samples is that Fe and Ti have been enriched from an external source. Deep sea nydrothermal vent waters may be a factor of 106 higher in Fe than average seawater (Von Damm et al. 1985). Palaganitization occurring in the vicinity of such hydrothermal activity could occur in solutions considerably higher in Fe (and Ti?) than average seawater. If the iron in these solutions is extracted to form Fe-oxide or Fe-hydroxide crystallites early during palagonitization and is not renewed, the first- formed palagonite will be higher in iron than later-formed palagonite. This would explain the relatively higher iron palagonite toward the outer portions of palagonite rinds in the high-Fe, low-Si deep sea dredge sample palagonite USNM 113521-69. 179 There is no general conclusion that can be drawn concerning compositional gradients in deep sea drill core palagonite rinds. Sample OOP 49 41QA 2-2 exhibits palagcnite rind compositional gradients similar to the trends seen in USNM 113521-69, with high iron enrichment toward the outermost (first-formed) portion of the rinds. As HicnigeaH above, this nay be indicative of the effect of Fe-rich hydrothermal solutions. Other drill core samples in which gradients were measured do not show such a pronounced high Fe effect. There are relatively high iron regions in OOP 37 335 10-4 palagonite rinds but the normalized FeO contents in the most iron-rich of these regions attains values of only about 25 wt.% compared to nearly 45 wt.% in ODP 49 410A 2-2 rinds. There is a general tendency toward decreasing Fe and slightly increasing Al concentrations toward the glass/palagonite interface in ODP 37 335 10-4 in the innermost 100 microns of the rinds. CaO and Mgo also appear to increase inward in rinds in this sample, perhaps indicating that the gradients are mainly due to a decreasing proportion of an Fe-rich component. ODP 37 332B 6- 2 has the lowest overall FeO and AI2O3 contents of any drill core palagonite in which Q4PA rind traverses were performed. Iron and calcium tend to remain constant across the rinds in this sample while aluminum decreases steadily toward the inner (later-formed) portions of the palagonite rinds. MgO Increases systematically toward the inner portions of the same rinds. There is no indication that there is an increasing proportion of an Mg-smectite (saponite) component which would correspond to this Increase in Mg and decrease in Al as total tetrahedral (IV)/total octahedral (VI) cation ratios remain close to 2.0 across the rind. More trioctahedral cations would shift the IV/VI values down toward 1.33. The increase of Mg may, therefore, be due to sorption of Mg2+. Compared to other drill core sample palagonites, ODP 37 180 332B 6-2 has the highest overall MgO content and is the only analyzed drill core sample which exhibits HO2 depletion in palagonite relative to glass. Note, however that there is no such correlation between Mg and Ti in the titanium depleted, lowHFe, high-Si dredge sample palagonites. 4.2 SEO0NDAKY MINERAR MJIHIGEHESIS. SOIUTION CONCTIlTRftTIONS. AND MASS BAIANCE A number of authigenic minerals are associated with palagonitized glass. Clay minerals, zeolites, and calcite are the most cannon, generally occurring as oanentaticm-authigenic minerals. In order to understand authigenesis in palagonitized deposits, the following observations must be addressed; 1) sane samples are zeolitized and some are not, some non-zeolitized samples are highly palagonitized; 2) zeolites are only seen as major alteration products associated with palagonite rinds that are aluminum depleted relative to parent glass; 3) phillipsite is not cannon in seawater altered samples fron Iceland but is the only zeolite seen in the zeolitized deep sea samples; 4) chabazite is, by far, the most (.xiiiiiun zeolite in fresh water altered samples and in Icelandic seawater altered samples; 5) clays are seen in many samples and their presence or absence does not appear to be related to palagonite composition. The occurrence of zeolite or clay is related to solution pH and the resulting speciation of dissolved aluminum (DeJong, Schramm, and Parziale 1983). Table 22 shows the relationship between hydrolysis number (nominally 181 Table 22. lelationship between pH and aluainua speciation (after OeJong et at., 1983) Hydrolysis NUMber1 Al Species 3 VI tow pH <1.3 Al(H2O)6 * contains Al 7 IV VI 1.5 to 2.5 A113 >2.6 (MgO present only) AlVI IV high pH 5 to 15 Al(0H)4* contains Al 1 Nominally bound ON/Al. 182 bound OH/M.) and aluminum coordination in solution. In the presence of silica, clay minerals will be favored at low pH as Al^1 dominates in solution. At high pH, Al^V dominates in solution and zeolites are precipitated. The rrence of clay cement prior to zeolites in pore spaces of the hyaloclastites from Kopavogur and Arnarnesvogur and in zeolitized vesicles in Tungufell pillow rim samples can, therefore, be interpreted as representing the solution pH increase expected from glass hydrolysis in closed systems of relatively low water/rock ratios. Whether the nontronite-like palagonite itself is restricted to lower pH conditions is more difficult to say. As previously mentioned, gels (like palagonite) form metastably in the conditions of high supersaturation with respect to more stable and better crystallized phases at the glass surface and, as such, may retain some of the original polymerization of the glass. As mentioned previously, abundant Mg in solution may cause aluminate clusters to revert to octahedral coordination. The clay mineral that has formed beneath zeolite in samples from Kdpavcgur and Tungufell is Mg-anectite (saponite). Crystallization of this clay lowers the Mg/Ca ratio in solution which is conducive to zeolitization. In a closed system, K*" can be removed from solution by adsorption or ion exchange on palagonite, thus the zeolites which form tend to be Ca-rich (chabazite). The aluminum content of palagonite appears to correlate with zeolitization in the sense that zeolites are not found or are rare where aluminum has not been depleted in palagonite relative to the glass precursor. The fact that aluminum has been depleted from a palagonite rind does not, however, imply that zeolitization will necessarily occur. Figure 82 183 Mg+(Mn) Palagonite Iceland Dr«dg« Samples octahedral cations Drill Cort Samplts Fe+(Ti) Figur» 82. Octahedral cation plot of palagonites assuming 22 oxygtna. An "*w" above a data point indicatea that it raprtatnta a seewater altered sample from Iceland. Filled symbols repreaent zeolitized aanplca. The shaded «rea includes all fresh water altered samples. 184 correlates zeolite occurrence with octahedral cation relationships for palagonites. Deep sea dredge samples with the lowest Al^1 content are not zeolitized. Ihe occurrence of zeolites is not clearly predictable based on the proportion of octahedral aluminum in palagonite. Aluminum loss during palagonitization is a neccessary but not sufficient condition for zeolite formation. Figure 83 shows the relationship between the AI2O3 content in palagonite (relative to glass) and rind thickness. Zeolites are not seen in samples where rind thicknesses are below 20 microns. However, in the auite of samples included in tinis study, there are no aluminum depleted palagonites that have rind thicknesses below 20 microns. Perhaps this relates to the higher dissolution rates expected under the high pH conditions which could produce aluminum losses. The occurrence of zeolites must relate to the effective solution volume (pore volume + flow rate * true exposure age) relative to the volume of glass reacted. If flow rates are low, the system may be considered closed. As an approach to a quantitative description of solution concentrations in natural samples, calculations were performed assuming no flow rate and initially deionized water. If palagonitization is an isovolumetric process then the volume of glass reacted is equal to the volume of palagonite. Hie volume relationships between palagonite and solution volume are -obtainable from thin section point counts (for pore volume) or vesicle size and rind thickness measurements (for vesicle relations). The amount of material released can then be expressed as total element concentrations either in solution or solid phases. Any flow rate greater than zero will lower the true dissolved element concentrations. Concentrations calculated for vesicles will be higher than true concentrations because of the limitations of thin section measurement 185 10 * o m net zaolftfztd :;:;:::::::::: • wolitootion field C o 1 o. 0 gf ; ? O •:f::ivXvX;:vXvX;:x < < SI ||| i || 10 10 10* 10' 10* Rind Thickness (microns) Figure 83. Rtlativ* Alj^S 'n P*l*flonite given T-2 Inferred Concentration* in Vesicles vesides with zeoil' ts 40.00 ; - 37 335 10-4 * • V«tlCl«« with e:oy 4 030.00 I e o • £20.00 i = «-i Tr-~1 C/1 • : - K0.-1 •• 4 • 10.00 * • 4 4 - 4 * * " 4 * ; * «4 *4* ** * 0.00 -< #*• f 0.00 200.00 400.00 * 600.0i 0 800.00 1CCC.00 Apparent Vesicle Diameter (microns) T-2 Inferred Concentration» in Vesicles 20.00 * • O • vMiCltl »iln ZtOlitll • • vtficict miVt eioy 1 - 37 333 10-4 15.00 • • V) ; • • a "3 10.00 : . • E • • • • - K0.-1 » 5.00 3 - «-i ... * . * § é 4 0.00 0.00 ZOO'.OO 400JO SOO'.OO 600.00 1000.00 Apparent Vesicle Diameter (microns) Figure 84. tnftrrvd total dissolved element concentrations of S' and Al in vesicles from sample T-2, Tungufell, Iceland. The plotted values assume that ill material lott as a result of palagonitfzation are put in solution and that solution volume is equal to vesicle volume (no flow rate). Concentrations shown for OOP 37 33S 10- 4, Ko.-I, R-1 Reyk.-I, and Tr.-1 are the inferred concentrations calculated for pore spaces in these samples. 188 concentration of Al vs. vesicle size relationships are shown for sample Reyk.- 1 in Figure 85. This figure indicates that concentrations were higher in smaller vesicles, an effect which is at least partly due to the geometric effects of thin section measurements as mentioned above. Palagonite is followed by cementation clay in the alteration paragenetic sequence of all clay-bearing samples. Clay overgrowing zeolites was not observed. The occurrence of zeolites is related to the precipitation of Mg- rich clay prior to zeolitizaticn, and either during or after palagonitization. Ca-rich chabazite preceeds Na-rich analcime in Kopavogur, Iceland sample Ko.- 1. Both zeolites overgrow calcite. The general sequence from calcium-rich zeolites (usually proceeded by K-rich phillipsite) to sodium-rich zeolite has been described at a number of locations (e.g. Iijima and Harada 1969; Hay and Iijima 1968a 1968b; Honnorez 1978; Brey and Schmincke 1980). As calcite preceeds zeolite in Kb.-l, high initial PQCE can be inferred. Externally buffered solutions (open system conditions) would not be expected to produce different zeolites with time if elemental proportions released from the glass were also constant with time. In closed systems, zeolite type will change with time as the solution responds to element release during palagonitization and element consumption during precipitation. Higher initial Si/ (Al+Fe) would favor chabazite over phillipsite, particularly if K*" is available. As the level of dissolved calcium in solution is lowered as a result of chabazite precipitation (and Si and Al are still available), sodium-rich varieties, such as analcime, are precipitated. System mass balance can be determined if phase compositions, densities, and volumes are known. If large amounts of material input are required to 189 Reyk.-1 Inferred Concentrations in Veaiotes 20.00 0.00 0.00 200.00 400.00 600.00 800.00 1000.00 12ÖÖ.0O ViÖÖ.'6Ö' 1600.00 1800.00 Apparent Vesicle Diameter (microns) Figure 85. Inferred total dissolved »lament concentrations of Al in vesicles from sample Reyk.*1, Reykjenes, Iceland. T< plotted values assume that all material lost ss a result of palagonifixation are put in solution and that solution volume is equal to vesicle volume (no flow rate). Concentrations shown for OOP 37 335 10- 4, Ko.-1, R-1 Reyk.-1, and Tr.-1 are the inferred concentrations calculated for pore spaces in these samples. 190 generate the observed alteration phases, then open system conditions are indicated. Similarly, if large amounts of material are lost free the system, then high flow rates may be inferred (high effective solution volumes may not allow solubility limits to be reached). Mass balance calculations were performed assuming that palagonitization was isovolumetric. A second calculation was performed assuming that Ti was conserved during palagonitization. Figure 86 shows the results for Iceland subglacial-volcanic sample Tr. -1 (Trolladyngia). Assuming isovolumetric palagonitization results in apparent enrichments in Ti, Al, Fe, and P above the amount originally present in the glass. Si, Mn, Mg, Ca, Na, and K are all lost from the system. This sample is cemented by a silica bearing Fe-oxide or hydroxide. The cement does not account volumetrically for much of the alteration products and, thus, does not greatly affect the mass balance except for Fe and Mn. There is no petrographic evidence for shrinkage during palagonitization (Figure 4) as would be required for Ti conservation. Calculation of the mass balance assuming conservation of Ti, however, indicates that Si loss would be to approximately 30% of the amount in the dissolved glass and than aluminum would then be depleted. All element concentrations in the alteration products could be accounted for in the amount originally in the glass. Net mass loss of Si, Mn, Ca, Na, and K would have been quite high. A 26 percent volume decrease from reacted glass to palaganite is required for Ti conservation. Ti conservation could also be obtained by decreasing the density difference between glass and palagonite by 26%. The mass balance in Icelandic subglacial-volcanic sample R-l is similarly dominated by palagonite as the clay cement is not abundant (Figure 87). In this sample, Si, Al, Fe, Mn, Mg, Ca, Na, K, and P all show significant net Figure 86. Mass balance results for fresh water altered sample Tr.-1 from Trolladyngis, Iceland. Amounts of elements in alteration products arc givtn in paretnt ralatfva to tht amount of dissolved glass. The upper diagram indieatts results assuming isovolumetrie palagonitization. The lower diagram assumes that all Ti is conserved in the palagonitization process. 100X totals indicate that the entire amount of that element involved in glass hydrolysis is used to make alteration products. Any amounts below 100X indicate system loss. Totals above 100% indicate that externally derived material is required to make alteration products. 192 R—1 Mass Balance (itovolumvtrie palogonitiiotion) 120 no - 100 - 90 - 80 - 70 - 60 - 50 t 40 1 77Å 30 77, 20 - 10 - O Si Ti Al Fe Mn Mg Cc No K P 17*71 Pol ftepl. Cloy RZ3 C«m Cloy K^ Z«oi R—1 Mass Balance (TI eenatont during polagonitization) 120 EZ3 Ffgurt 67. HMS bilanea rttutti for fr«th wattr altartd tampla R-1 from Raudafall, Ictiind. Anounta of alafltants in alttration products art gfvtn in pareant ralativa to the amount of dissolved glass. The upper diagram indicates results assuming isovolumetric palagonitization. The lower diagram assumes that all Ti is conserved in the palagonitization process. 100X totals Indicate that the entire amount of that element involved in glass hydrolysis is used to make alteration products. Any amounts below 100X indicate system lost. Totals above 100X indicate that externally derived material is required to make alteration products. 193 mass loss to solution (and out of the system). The concentration of Ti in palagonite is quite close to the expected value with the assumed densities and isovolumetric reaction with no enrichment. Clays are abundant in Iceland seawater-altered sample Reyk.-l from Reykjanes. The presence of this iron-rich clay indicates that significant amounts of Fe must be brought into the system in order to form all alteration products with their indicated compositions (Figure 88). Isovolumetric reaction would lead to the conclusion that Ti, Mn, Ca, and P have been lost from the system and that Si, Al, Fe, Mg, Na, and K (note the potassium value plotted is .10K) have all been enriched above the amount made available by glass dissolution. The assumption of conservation of Ti would inply that either there has been volume increase during palagonitization or that palagonite is actually more dense than glass. The observed apparent Ti loss may, more likely, be real. Note that a significant amount of Ti is incorporated into the clay. By revising the assumed clay density upward slightly, the calculated loss of Ti from the system could be eliminated. A source for Ti in the clay is neccessary, the most likely possibility being the release of Ti into solution during palagonitization. The mass balance of sample Kb.-l from Kbpavogur, Iceland is shown in Figure 89. Isovolumetric palagonitization implies that Ti and K are enriched above the amount of these elements in the reacted glass (note that the K values shown are in units of .10K). Si, Al, and Na are nearly conserved in the system after zeolitization, and no externally derived materials would be required to form all alteration phases with their indicated compositions except for Ti and K. Calculating mass balance assuming conservation of Ti 194 Reyk. —1 Mass Balance (l*evolum«tric poiaqonitizotion) 2 I F« Mn Mg Co .10K GT71 Poi *«pt. Cloy V77X Cam. Cloy Z«oi Reyk. —1 Mass Balance (TI eenttont during eofogonitizotion) | 0. Si Tl Al r« Mn Mg Co No .10K E3 •« K~^ R«pl. Cloy C«m. Clay Zaol Figur* 88. Mais balanc* rtaultt for itawattr a(t«r«d »mpt« R«yk.-1 from Raykjant», Iceland. Aneuntc of alamanta In altaration products ira givan in parcant ralative to tha amount of diasolvtd glaas. Tha upp*r diagram indicata» raaultt aatuming isovoluwatrfc palagonitizati'on. Tha lowar diagram aiaumat that all Ti i> conaarvad in tha palagonitizatlon procaaa. 100X totals indicata that tha entire amount of that * lament involved in glass hydrolysis is used to make alteration products. Any amounts betow 100X indicate system loss. Totala above 100% indicata that externally derived material is required to make alteration products. 195 Ko. —1 Mass Balance (itovolumrt. ic pelogonitizotion) 14O i o 2 PT1 Pol Ko. —1 Mass Balance (Ti constant during oofOQonititotion) Figurt 89. Man balanct rttultt for tttwatir alttrad sanplt Ko.-I from Kopavogur, Icttand. Aoounta o? tlttMntt in altaration products ara givan in pareant ralativt to tht aawunt of dissolved glatt. Tha uppar diagran indicatat ratultt tttuming itovolumttric palagonitization. Tha lowtr diagram attumat that all Ti it contarvad in tha palagonitization procatt. 100X totalt indicata that tha tntira amount of that alamant involvad in glatt hydrolytit it utad to maka altaration products. Any amount! balow 100X indicata tyttam lott. Totalt above 100X fndicata that axtarnally darivad matarial it raquirad to make altaration productt. 196 during palagonitizatian indicates that all elements other than Ti and K would be lost from the system. Externally derived K is still required. A 20 percent decrease in volume from glass to palagonite (or 20 percent lower denstiy differenoe between glass and palagonite from the assumed densities) is required for Ti conservation. Isovolumetric palagonitization of glass in sample A-l shows that all elements would be lost to the local aqueous solution, including potassium, with the exception of titanium (Figure 90). Cementation by clay and zeolite does not require using all remaining amounts of other elements lost during palagonitization except for potassium, which is evidently enriched above the amount in glass at this point. The assumption of Ti conservation during palagonitization results in increased calculated mass losses for all elements. A volume decrease of 26 % from glass to palagonite is required for Ti conservation without enrichment from external sources during palagonitization. The apparent K depletion during palagonitization is inconsistent with seawater alteration for this sample even though the Amarnesvogur hyaloclastite deposits were collected from wave-eroded outcrops at the shoreline of the inlet. The net enrichment of K with cementation of the sample is typical of other seawater altered deposits. Interactions with fresh water or "lix*^ fresh water and seawater during the time of palagonitization could have caused the apparent depletion. The alteration history was undoubtedly complicated by changing sea level during the Pleistocene and also, perhaps, by fresh groundwater outflow through the deposits. Mass balance calculations were also done for sample ODP 37 335 10-4, one of the zeolite-bearing deep sea drill core samples. The calculation was done 197 A—1 Mass Balance (iaovolumatric pologonitizotion) 2 I f F71 Pal A—1 Mass Balance (Ti conrtorrt during pologonitization) % 1 Si Ti Al F* Mn Mg Co No Pal . Clay C«m. Clay Z«el Figurt 90. Nats balsnet rttult* for »tawattr alttrtd lampla A-1 fro* Arnarntavogur, Iceland. Amount! of alaiMnt* in alttration products ara givan in pareant raiativa to the anount of disaolvad glas*. Tha uppar diagram indicate» results assuming fsovolumatric palagonitization. Tha tower diagram assumes that all Ti is consarvad in tha palegonitization process. 100X totals indicate that the entire amount of that element involved in glass hydrolysis is used to make alteration products. Any amounts below 100X indicate system loss. Totals above 100X indicate that externally derived material is required to m**t alteration products. 198 using the volume relations between palagonite and zeolite in a centimeter long fracture in a glass fragment which is part of a glassy breccia. The zeolite which fills the fracture is low Ca-phillipsite. Calculating mass balance assuming isovolumetric palagonitization (Figure 91) shows that Ti and Fe are only slightly enriched above the amount originally in the glass. Potassium is highly enriched and is abundant in both palagonite and zeolite (note, again, that the K values presented are in units of .10K). With zeolitization, Na is approximately conserved in the system. Si, Al, Mh, Mg, Ca, and P all show net mass loss from the system. Note that the apparent losses in Mg and Ca are much greater than was the case for the Icelandic seawater altered sample Ko.-l which is also zeolitized. Because there is only a slight apparent increase in Ti concentration during palagonitization if isovolumetric reaction is assumed, recalculating assuming conservation of Ti does not change the mess balance much in ODP 37 335 10-4. A volume decrease of 4 % would account for all the titanium present in palagonite. Of the four seawater-altered samples for which mass balance calculatioiis were performed (Reyk.-l, Ko.-l, A-l, and ODP 37 335 10-4), only the non- zeolitized sample Reyk.-l shows significant mass increase for Si, Fe, and Mg. This sample also shows net mass loss of Ti. Open system alteration conditions and high flow rates are thus indicated for Reyk.-l, similar to conditions inferred for the clay-filled vesicles in Tungufell samples (see Figure 81). Loss of Si and Al to the soluticn were high during palagcnitization and K, Na, and Ca should have been readily available in solution from both glass dissolution and external sources. These elements are the major components of zeolites. Solution concentrations, as discussal above, may not have been high enough for zeolite precipitation. Local solutions must have been magnesium 199 ODP 37 335 10-4 Mass Balance (i»ovoium#tric polagonitizotion) Mn Mg Co No . 10K P Cam. Cloy (^3 Z«ol ODP 37 335 10-4 Mass Balance (T1 constant during poloqonitizotion) 200 Ti Al Fe Mn Mg Co No -10K P R«ol. Cloy C«m. Clay Figurt 91. Matt balanct rttultt for d«tp ••• drill cor* tampl* OOP 37 335 10-4. Amounts of tltiMntt in alttration products irt given in ptrctnt rtlatfv* to the amount of dissolved glass. Th« uppar diagraM indicatts results assuming isovolumetric palagonftization. Th* lower diagram sssumes that all Ti is conserved in the pa Iagonifixation process. 100% totals indicate that the entire amount of that element involved in glass hydrolysis is used to make alteration products. Any amounts btlow 100X indicate system loss. Totals above 100X indicate that externally derived material is required to make alteration products. 200 rich as a net mass input of this element is indicated (Figure 88). High magnesium may suppress zeolitization and allow precipitation of clays. Despite complete cementation of the fracture in deep sea drill core sample ODP 37 335 10-4, significant mass losses of Si, Al, Vq, and Ca have still occurred (slightly over 20 % loss in Si and Al, 70 % loss of Mg, and 95 % loss of Ca. With such high Mg losses to the solution, it is unclear why smectite was not precipitated rather than phillipsite. Availability of Fe in solution may be neccessary for clay formation, an was apparently the case in sample Ko.-l. 4.3 Palagonitizatian rates cannot be directly inferred from rind thickness measurements and sample ages. This is because the actual exposure age of the sample, that is, the time the glass has actually been in contact with water, is not known. True exposure age will always be somewhat less than the sample age so that rates inferred using sample ages (apparent rates) will be minimum values. The apparent rates thus inferred for samples from Iceland are given in Table 23, those for deep sea dredge samples in Table 24, and those for deep sea drill core samples are given in Table 25. The rind thickness/age relationships for each sample are plotted in Figure 92. Definitive statements concerning the maximum alteration rates of glasses are not possible because of the difficulties in assessing the actual contact times. However, if continuous contact of glass with solution throughout the 201 Table 23. Palagonitization rates inferred for Iceland samples from apparent rind thickness and corresponding sample age. Rate (microns/IOOOvr.) Sample Minimum Maximum Iceland Fresh Water Tr.-1 .007 1.88 M-1 .014 10.0 M-7 .029 5.00 Mid.-1 0.0 5.00 Mid.-2 0.0 5.00 Mid.-3 .007 2.50 R-1 .021 3.75 R-2 .004 .625 BI.-1 .029 3.75 SI.-4 .029 3.75 T-2 .013 .257 T-3 .006 .286 T-4 .002 .043 V-1 .004 1.25 Iceland Seawater Br.-I .021 3.13 Br.-2 .029 2.50 K0.-1 .021 3.75 A-1 .057 7.50 Reyk.-1 .029 10.0 Reyk.-2 .029 10.0 Rate based on maximum age. 2Rate oated on minimum age. 202 Table 24. Palagonitization rates inferred for dredge samples from apparent rind thickness tnd corresponding sample age. Rate (microns/IOOOvr.: Sample Minimum1 Maximum2 Reykjanes Ridge USNM 113521-27 .027 .600 USNM 113521-69 .001 .100 USNH 113531-6 .007 ?5 USNM 113427-25 .003 ? Hid Atlantic Ridge USNM 113487-3 .109 ? USNM 113715 .010 ? USNM 111308 .0003 ? USNM 110728-4A" .001 ? Galapagos Ridge Area USNM 113152-17 .001 7 USNM 113434-33 .001 ? USNM 113434-3* .002 ? USNM 113434-15 .003 ? USNM 113249-705 .050 7 USNM 114045-3A .025 .250 florda Ridge Area USNM 111238-1 .0002 ? Rate based on maximum age. Rate based on minimum age. Interior fractures. Outer surfaces. Question mark indicates that limiting surfaces cannot be defined. 203 Tablt 25. Palagonitization rates inferred for drill core samples from apparent rind thickness and corresponding sample age. Rate (microns/IOOQyr.) Sample Minimum1Maximum2 Rey it janes Ridge OOP 49 409 25-1 017 .025 ODP 49 410A 2-1 002 .005 OOP 49 410A 2-23 004 .012 OOP 49 410A 2-24 026 .253 Mid Atlantic Ridge ODP 37 332B 6-2 006 .023 OOP 37 332B 45-1 006 .023 OOP 37 334 16-5 316 ?5 ODP 37 334 17-13 001 .012 ODP 37 334 17-14 179 ? ODP 37 335 10-4 002 .015 Rate based on maximum ageg . Rate based on minimum age. 'interior fractures. surfaces. 'Question mark indicates that limiting surfaces cannot be defined. 204 10 S Iceland Dredge Samples Drill Core Samples 1111\ I I I I 11 III T i ' I 10" 10 "1 1 10 102 Age (m.y.) Figure 92. Rind thfcknesa va. age for samples from Iceland, deep sea dredge samples, and deep sea drill core tanptts. The diagonal lints are ratea, with indicated values given in microns/1000 yr. 205 history of the sample is ?>p- Grambow et al. (1985) attribute some of the difference in apparent rates to the difference between k+, the forward rate of reaction, and k, the final rate. This difference between k+ and k is due to the effects of silica saturation, which controls the final rate of reaction. The forward rate was determined to be on the order of 3 to 20 microns/1000 yr. The final rate appears to be around .1 microns/1000 yr. The values for the forward rate were obtained from samples from British Columbia and from the literature (Bryan and Moore 1977). Most of the samples included in this study have apparent rates near or below .1 micron/1000 yr. Samples altered in the deep sea environment, particularly dredge samples, have had the greatest chance of being in contact with water throughout their history. Even so, the may-imm apparent rates for those dredge samples for which minimum ages can be constrained (. 6 microns/1000 yr. for USNM 113521-27; .1 microns/1000 yr. for USNM 113521-69; .25 microns/1000 yr. for USNM 114045-3A) do not approach the palagonitization rates reported for other open ocean samples, such as: 2.6 to 4.3 microns/1000 yr. (Hekinian and Hoffert 1975); 3 to 30 microns/1000 yr. (Moore 1966); and 15 microns/1000 yr. at Surtsey (Jakobsson and Moore 1986). The wide disparity in rates for deep sea drill core samples may be attributable to the effects of sealing by authigenic cement. 206 5 O0NCLDSXGNS Laboratory experiments have demonstrated that basaltic glass and borosilicate nuclear waste-form glass are similar in terms of short-term corrosion processes (Malow, Iutze, and Ewing 1984; Lutze et al. 1985; Byers, Jercinovic, and Ewing 1986). Modelling the long-term alteration of borosilicate glass requires a means of verification which may be approached through the study of natural analogues, specifically, naturally altered basaltic glasses (Ewing and Jercinovic 1987). The information obtained from naturally altered basaltic glasses which is pertinent to the evalutation of the corrosion of nuclear waste-form glass is as follows: 5.1 CORROSION MECHANISM The alteration of basaltic glass proceeds by hydrolytic reconstruction, forming gel-like surface layers (palagonite) which replace the glass. Evidence of glass alteration exists in every sample studied. Where at least some fresh glass remains, the glass surface is dissolution pitted and a surface layer has formed. Palagonite preserves the original glass form throughout alteration. With the possible exception of the high-Fe, low-Si deep sea dredge samples, the replacement reaction appears to be isovolumetric. Microcrystals within the glass (olivine, pyroxene, plagioclase) remain unaltered, even when completely surrounded by palagonitized glass. Some unaltered glass is seen in samples as old as 13 m.y. (ODP 37 335 10-4); whereas, completely altered glass can be found in younger samples (USNM 207 113487-3, 9.5 to 11.0 m.y.). There does not appear to be any consistent maximum rind thickness which denotes the point at which no further alteration occurs, thus palagonitization is not diffusion-limited until deposits are sealed by authigenic cements. Glass dissolution must, therefore, be surface- reaction controlled. Because palagonitization is non-isochemical, the elemental release to solution is incongruent. 5.2 ALTERATION PRODUCTS Palagonite, the gel-like (colloidal) replacement product of basaltic glass, is seen in all samples studied. AIM steadies reveal that sane crystalline structure is present in the palagonite and that the degree of crystallinity can vary widely within a rind on a scale of tens of nanometers. Crystallite sizes do not exceed 50 nm in size. AIM also shows that palagonite rinds may be comprised of several separate phases. In the case of high-Fe, low-Si deep sea dredge palagcnites, both clay (smectite) and Fe-oxide are formed. The colloidal nature of palagonite implies precipitation under conditions of high supersaturation, forming essentially as critical nuclei with limited crystal growth. Palagonitization is, therefore, a non- equilibrium process with regard to more stable (i.e., better crystallized) All palagonites analyzed exhibit losses in silica relative to the amount present in precursor glass. Iron and titanium are almost always retained or enriched in the palagonite rind relative to parent glass. Some deep sea dredge palagonites appear to be highly iron-enriched. No palagonite shows 208 large iron depletions. The observed Fe concentrations may be a result of Fe enrichment from an external source ("hydrotheraal" water), rind shrinkage during palagonitization, or lew density of the palagonite. Apparent Ti depletion is oauuoji in sane low-Fe, high-Si deep sea dredge samples. Clearly, even in the simplest environment (seawater, open ocean), significant variability in palagonite composition occurs. Significant difference also exist among seawater-altered samples from Iceland. Palagonites from Kopavogur, Arnarnesvogur, and Reykjanes are Al-depleted relative to parent glass while Al is retained during palagonitization of samples from Brimnes. In Icelandic subglacial palagonites, depletion of alkaline earths is commonly 10% of the original concentration in glass and alkalis are nearly completely depleted. Seawater-altered samples have palagonite rinds that are highly enriched in K, and Na is retained in the rind to a much greater extent than for fresh-water formed palagonites. This is the one major, consistent difference between seawater and fresh-water palagonitization. Some palagonite rinds exhibit compositional gradients across the rind, most commonly, with lower Al-content in the outermost (first-formed) portion of the rind, changing to higher Al in the inner (later-formed) portions of the rind. This increasing Al-content is generally accompanied by a corresponding decrease in Fe. Mg is generally more abundant in the outermost portions of palagonite rinds, becoming progressively less concentrated toward the inner portions of the rinds. Ca inreases correspondingly. These trends appear to reflect changes in solution composition with time. The compositional gradients are more strongly developed in rinds associated with zeolite cement. Closed-system alteration is indicated for such samples. 209 Glass hydrolysis provides elements for authigenic reactions during and/or after palagcnitization. Zeolites are abundant cements in the palagonites which show significant loss of AI2O3 relative to the glass progenitor. The Al-depletion nay be due to high pH as rinds are generally thicker (higher reaction rate) where zeolites occur. Seme Al-depleted palagonites are not associated with zeolites, a result of high effective solution volumes. Mg- rich smectite clay (saponite) is usually precipitated prior to zeolite. This lowers the Mg/Ca ratio in solution. If Mg is released during palagonitization and some Fe is avialable in solution, smectites will precipitate instead of zeolite. The behavior of Mg during palagonitization is critical in determining the subsequent alteration-paragenetic sequence. Chabazite is, by far, the most cannon zeolite in both fresh water- altered and seawater-altered samples from Iceland. Analcime is found in one sample from Iceland (Kopavogur), overgrowing chabazite. Low-Ca phillipsite is the only zeolite identified in the deep sea samples of this study. 5.3 MASS BMANCE Mass balance calculations indicate that truly closed-system alteration did not occur in any of the deposits sampled. Some mass loss always occurs, particularly of alkaline-earths. In fresh water-altered samples, Si, Mg, Ca, Na, and K are lost from the system. Fe and Ti are often enriched in clay- bearing, fresh water-altered samples, indicating net mass input. Nearly closed system alteration has occurred in zeolitized samples from Iceland (Tungufell, Kopavogur, and Amarnesvogur). In these cases, Si, Al, Mg, Ca, 210 and Na, released into solution during palagonitization, are used to form authigenic minerals (saponite and zeolites), thus resulting in limited net mass loss. Al loss during palagonitization and zeolite crystallization result from high pH and low effective solution volume (low pore volume and low flow rate). The mass balance for seawater-altered samples indicates that the K content in the alteration products results primarily from uptake of K frön seawater. 5.4 ftLTEF^TIQN RATES "Apparent" palagonitization rates are from .001 to 1 micron/1000 yr. All rates inferred frcm palagonite rind thickness and age are minima rates because the actual time the glass is exposed to the solution is less than the age of the sample. Maximum rates nay possibly be obtained from uncemented deep sea dredge samples. The highest rate inferred for an uncemented dredge sample is .6 microns/1000 yr. Palagonitization rates as high as 30 microns/1000 yr have been reported for other dredged glasses (Moore 1966; Moore, Fornari, and Clague 1985). Dissimilarities in rind thickness for sanples of the same age can be explained by two effects: 1} differences in solution corapositian either initially or with time resulting in different reaction rates (i.e., large effective solution volumes, a result of high initial volume relative to surface area or high flow rates, will be less effected by elemental release during palagonitization than smaller effective solution volumes); 2) differences in true exposure age of glass to solution compared to the age 211 of the sample, this difference say be due to either intermittent exposure (e.g., alteration in unsaturated zones of deposits) or sealing by authlgenic minerals. 212 The UNM natural analogue program was initially supported by the U. S. Nuclear Regulatory Ocmnissicn, through Argonne National Laboratory (Contract No. 60-84-254, FIN No. A2254-4 to R. C. Ewing). The deep sea dredge and drill core sample collection was assembled, and preliminary data was gathered for Icelandic samples and deep sea dredge samples under the scope of the U. S. NRC project. W. G. Melson and T. O'Hearn at the U. S. National Museum of Natural History, and R. C. Hayman of the Ocean Drilling Program are acknowledged for their help in obtaining research samples. Sveinn Jakobsson and Hjalti Franzscn provided very helpful discussions and direction prior to field wor> in Iceland. The analytical electron microscopy results were obtained by Takashi Murakami (on leave fran the Japan Atonic Energy Research Institute) at UNM. Analytical work was canpleted in the Electron Microbean Analysis Facility of the Department of Geology, UNM supported by NSF, NASA, DOE-BES and the State of New Mexico. 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