Clinoptilolite series
Clinoptilolite-K |(K,Na,Ca0.5,Sr0.5,Ba0.5,Mg0.5)6(H2O)20|[Al6 Si30 O72]
Clinoptilolite-Na |(Na,K,Ca0.5,Sr0.5,Ba0.5,Mg0.5)6(H2O)20|[Al6 Si30 O72]
Clinoptilolite-Ca |( Ca0.5,Na,K, Sr0.5,Ba0.5,Mg0.5)6(H2O)20|[Al6 Si30 O72]
Morphology:
Monoclinic 2/m, platy crystals with prominent {010} face, modified by and {001}, {100}, {111}, {201} and {110}.
Physical properties:
Cleavage: {010} perfect.
Hardness: 3.5 - 4
D: 2.14 – 2.17 gm/cm3.
Luster: vitreous, pearly on {010}. Streak: white.
Clinoptilolite-Ca with smectite on surface of rhyolite ignimbrite, Richardson’s Ranch north of Madras, Jefferson County, Oregon, USA. Width of view 5 mm.
Optical properties:
Color:
Colorless, white, yellowish, pinkish, orange to red; colorless in thin section
Clinoptilolite-Na partial pseudomorph of glass shard (0.25 mm in length) from the Miocene, Barstow formation, San Bernardino County, California, USA. Crossed polars. Image courtesy of R. A. Sheppard (see Sheppard and Gude 1969).
Biaxial (+ or -),
Clinoptilolite-K:
α 1.476 - 1.478, β 1.477 - 1.479 , γ 1.479 - 1.481 , δ 0.003, 2Vz 58 - 73°, Y = b, Z ˄ c 38° - 58° (or Z = b).
Dispersion: r > v, distinct, crossed
Clinoptilolite-Na:
α 1.474 - 1.478, β 1.475 - 1.479, γ 1.478 - 1.481, δ 0.003 – 0.004,
2Vz 32 - 95°, X = b, Z ˄ c 18° - 23° (or Z = b).
Dispersion: r < v, moderate
Clinoptilolite-Ca:
α 1.481, β 1.484, γ 1.488, δ 0.007, 2Vz 70°, Z = b, X ˄ c 20°.
Dispersion: r < v, distinct
Crystallography:
Unit cells:
Clinoptilolite-K
a 17.688(16), b 17.902(9), c 7.409 (7) Å, β 116.50(7)°.
Z = 1, Space group C2/m, C2, or Cm.
(Ogihara and Iijima 1990, off-shore borehole, Japan)
Clinoptilolite-Na
a 17.627(4), b 17.955(4), c 7.399(4) Å, β 116.29(2)°.
Z = 1, Space group C2/m, C2, or Cm.
(Sheppard and Gude 1969, Barstow Formation, San Bernardino County, California, USA)
Clinoptilolite-Ca
a 17.660(4), b 17.963(5), c 7.400(3) Å, β 116.47(3)°.
Z = 1, Space group C2/m, C2, or Cm.
(Koyama and Takéuchi 1977, Kuruma Pass, Fukushima Prefecture, Japan)
Names:
The origin and use of the name clinoptilolite has a convoluted history. The type locality is
on the ridge trending several kilometers northeast of Hoodoo Peak just outside the eastern boundary of Yellowstone National Park, Park County, Wyoming, USA, where the mineral occurs cavities in decomposed basalt breccia. The crystals were first identified by Pirsson (1890) as mordenite, based solely on chemical analysis. The platy habit and optical properties led Schaller (1923, 1932) to conclude that it was a monoclinic dimorph of ptilolite, a finely fibrous mineral described and named by Cross and Eakins (1886) for material occurring at Table Mountain, Colorado, USA and later shown to be mordenite. Schaller, therefore, named these crystals clinoptilolite, even though the morphology was very similar to heulandite. Based on X-ray diffraction data Hey and Bannister (1934) determined that the Hoodoo Peak material and heulandite are isostructural and recommended that the term clinoptilolite not be used.
Mason and Sand (1960) proposed a new definition for clinoptilolite, based on alkali-dominant and high Si/Al compositions. Mumpton (1960) simultaneously suggested that the name be applied to those samples that remained stable following overnight heating to 350°C. This method works particularly well for the fined-grained replacements of vitric tuff. During the following years the compositional gap observed by Mason and Sand (1960) was filled by the analysis of many newly discovered samples. Even so, the subcommittee reviewing the nomenclature of the zeolite group (Coombs et al. 1997) chose to retain the two mineral names, and proposed to keep both the heulandite and clinoptilolite names and to separate them based on the framework composition at Si/Al = 4.0. For a discussion of this nomenclature problem and some guidance in distinguishing between heulandite and clinoptilolite, see Bish and Boak (2001).
Both names were also raised to series status to include several species based on the dominant cation content. The clinoptilolite series comprises three species. Clinoptilolite-K is the new name for the original material from the ridge east of Hoodoo Peak, Wyoming. Clinoptilolite-Nais a new name for Na dominant crystals with the suggested type example from the Barstow formation, San Bernardino County, California, USA, and clinoptilolite-Cafor Ca dominant samples with type examples from Kuruma Pass, Fukushima Prefecture, Japan.
Crystal structure:
Both clinoptilolite and heulandite possess the same tetrahedral framework (labeled HEU) and form a continuous compositional series sometimes referred to as the heulandite group zeolites. The crystal structures of clinoptilolite and heulandite are mostly described to be monoclinic, space group C2/m(e.g. Alberti 1975, Koyama and Takéuchi 1977, Bresciani-Pahor et al. 1980, Alberti and Vezzalini 1983, Hambley and Taylor 1984, Smyth et al. 1990, Armbruster and Gunter 1991, Armbruster 1993, Gunter et al. 1994, Cappelletti et al. 1999). However, lower symmetries such as Cm and C1 have also been reported (Alberti 1972, Merkle and Salughter 1968, Gunter et al. 1994, Yang and Armbruster 1996, Saniet al. 1999, Stolz et al. 2000a). The HEU framework contains three sets of intersecting channels all located in the (010) plane. Two of the channels are parallel to the c-axis: the A channels are formed by strongly compressed ten-membered rings (aperture 3.0 x 7.6 Å) and B channels are confined by eight-membered rings (aperture 3.3 x 4.6 Å). C channels are parallel to the a-axis, or [102] and are also formed by eight-membered rings (aperture 2.6 x 4.7 Å).
Alberti (1972) concluded that the true probable lower symmetry of heulandite cannot reliably be extracted from X-ray single crystal data because of strong correlations of C2/m pseudo-symmetry related sites during the least-squares procedure. Thus C1, C1, Cm, C2, C2/m are possible space groups for clinoptilolite and heulandite. Akizuki et al. (1999) determined by optical methods and X-ray diffraction that a macroscopic heulandite crystal is composed of growth sectors displaying triclinic and monoclinic symmetry where the triclinic sectors are explained by (Si,Al) ordering on the crystal faces. Yang and Armbruster (1996) and Stolz et al. (2000a,b) stated that, owing to correlation problems, symmetry lowering in heulandite can only be resolved from X-ray data when investigated in cation-exchanged samples where the distribution of non-framework cations also reflects the lower symmetry.
Differing degrees of (Si,Al) ordering over the five distinct tetrahedral sites (assuming C2/m space group) have been reported for both heulandite and clinoptilolite. In all refinements, the tetrahedron with the highest Al content, T2, joins the “sheets” of T10O20 groups by sharing their apical oxygens. A neutron diffraction study by Hambley and Taylor (1984) located the majority of the H atoms and found (Si,Al) ordering values similar to other C2/m refinements. Additional (Si,Al) ordering, due to lower symmetry (C1 or Cm), was resolved by Yang and Armbruster (1996), Sani et al. 1999, and Stolz et al. (2000a,b).
Two main channel cation sites have been reported by all researchers and at least two more sites of lower occupancy have been reported by others (e.g. Sugiyama and Takéuchi 1986, Armbruster and Gunter 1991, Armbruster 1993). These sites commonly contain Na, Ca, K, and Mg, with Na and K predominantly close to the intersection of the A and C channels and Ca located in the B channel. The Na site in the A channel generally also contains Ca, whereas the Ca site in the B channel is mostly Na free. K and Na occur in nearby sites, but K is more centered in the C channel. Both can be distinguished by their different distances from the framework. Na, K, and Ca ions are on the (010) mirror plane, present in the C2/m or Cm symmetry, and they coordinate to framework oxygens and channel H2O molecules. In one refinement, Na was nine-coordinated with to four framework oxygens and five strongly disordered and partially occupied H2O molecules, whereas both Ca and K were eight-coordinated to four framework oxygens and four channel H2O molecules (Gunter et al. 1994). Mg commonly resides in the center of the A channel, coordinated only to six disordered H2O molecules (Koyama and Takéuchi 1977, Sugiyama and Takéuchi 1986, Armbruster 1993).
Clinoptilolite and heulandite contain differing amounts of H2O as a function of their non-framework cation chemistry (Bish 1988, Yang and Armbruster 1996) and hydration state. The H2O molecules occurring in the B channel (coordinated to Ca) are commonly fully occupied, whereas those occurring in the A channel are generally only partially occupied (Koyama and Takéuchi 1977, Armbruster and Gunter 1991). The structural mechanism of dehydration and accompanying framework distortion were studied by Alberti (1973), Alberti and Vezzalini (1983), Armbruster and Gunter (1991), and Armbruster (1993).
The crystal structure of clinoptilolite-Na (Agoura, California) with cation positions from the refinement of Koyama and Takéuchi (1977). Typically clinoptilolite contains 4 to 7 cations per unit cell (Deer et al. 2004).
Chemical composition:
Representative analyses of the three clinoptilolite species are plotted in the figures below. These species occur as diagenetic alteration products of volcaniclastic sediment in saline lakes and marine accumulations, including pelagic and siliceous clays in deep sea sediments, and in cavities in volcanic rocks ranging in composition from basalt to rhyolite. By the new definitions (Coombs et al. 1997) these species have high Si contents, Si/Al greater than 4.0, and TSi greater than 0.80. Clinoptilolite-K is the most widespread, mostly because deep sea clinoptilolite is K-dominant. However, there are also many occurrences in rhyolitic tuffs from terrestrial and marine environments. Both clinoptilolite-Na and clinoptilolite-Ca occur in a wide range of environments, including diagenetic replacement of rhyolitic volcaniclastic rocks, active hydrothermal systems, and fractures and cavities in volcanic rocks. Mg occurs in almost all clinoptilolite, but high amounts (greater than 1 weight %) may reflect included smectitic clay in those from volcaniclastic rocks. As in other zeolites, Fe is most likely Fe3+ and resides in tetrahedral sites, but amounts over 0.5 atoms/cell may be from included hematite. Sr and Ba are much less common in clinoptilolite species than heulandite. Nonetheless, Mg, Sr, and Ba should always be sought when analyzing a member of the heulandite structural group.
R2+ - R+ - Si compositional plot (left) and Na - (Ca+Mg+Sr) - K plot (right) of the high quality clinoptilolite series analyses compiled in Deer et al., (2004). Black squares represent heulandite analyses, all with Si/Al less than 4.0. In the Na - (Ca+Mg+Sr) - K plot the solid circles represent compositions from terrestrial occurrences, while the open circles represent those from the deep sea.
Occurrences:
Clinoptilolite series minerals are by far the most common zeolites. These minerals occur in rocks and sediment seemingly formed in widely disparate environments, which include deep sea sediment, continental accumulations in thick basin fills and in shallow lakes, and in some lava flow sections. The following summary is adapted from Deer et al. (2004), who also discuss conditions and mechanisms of zeolite formation.
Diagenesis of sediment and sedimentary rocks.
Minerals of the clinoptilolite series occur in diverse, low temperature environments, mostly, but not exclusively, as a result of alteration of volcanic rocks and fragmental debris. This alteration occurs as a reaction between volcanic glass and interstitial solutions during diagenesis. The composition of the zeolite to form is dependent on the compositions of the glass and of the solution. Clinoptilolite series minerals commonly replace high silica rocks, those ranging from andesite to rhyolite. Because of their size, and hence of economic importance, the replacement of ignimbrite sheets (welded pyroclastic flow deposits) that may occur in shallow marine environments and in terrestrial basins are of particular importance.
Diagenesis of marine pyroclastic and volcaniclastic sequences. Aleksiev and Djourova (1975) and Yanev et al. (2006) show that Paleogene volcanism occurring in the northeast Rhodope Mountains of southern Bulgaria, produced extensive deposits of sub-aqueous pyroclastic flows, fall-out tuff, and volcaniclastic sands all of which accumulated in a shallow marine basins. Irregular replacement of these deposits suggests that suitable conditions for the growth of zeolite minerals were controlled by local geology and fluid sources. Zeolitic alteration occurred from solutions consisting of marine water heated by the anomalous geothermal gradient of the active volcanic areas or/and by the hot pyroclastic rocks. Although individual units are incompletely exposed, composite vertical sections have the generalized vertical diagenetic mineral zoning: analcime, clinoptilolite-Na, clinoptilolite-Ca, clinoptilolite-K, and uppermost unaltered glass.
The Castilla, Tasajeras, Las Pulgas, and Caimanes clinoptilolite deposits in the western region of Cuba appear to be this type (Gonzalez et al. 1991).
Diagenesis in hydrologically closed systems. In general alteration in hydrologically closed systems occurs where the glass particles of fall-out tuff react with interstitial lake water, incorporated at the time of deposition. For a zeolite mineral to be produced, this water must have high salinity and high pH, usually higher than those of sea water. This situation can occur where rhyolitic ash falls into saline and alkaline lakes that form in arid environments. Replacement of vitric tuff in hydrologically closed systems tends to produce beds of nearly pure zeolite that have substantial economic value. Tuff beds replaced by this mechanism have a characteristic lateral variation from unaltered glass at the lake margin through an intermediate zeolite zone to a potassium feldspar zone near the lake center. This type of alteration process was recognized early by Hay (1966), but the clear relationships at the Pleistocene Lake Tecopa, Inyo County, California, USA, (Sheppard and Gude 1968), provide the conceptual basis upon which other occurrences are compared.
The exposed lake deposits of Lake Tecopa cover an area of about 230 km2, and consist of 67 m of mudstone with interlayered vitric, rhyolitic tuff, comprising 8 to 12 percent of the section (Sheppard and Gude 1968). Tuff beds range in thickness from 4 cm to 4 m. Post Pleistocene erosion has exposed the lake beds in Badlands-type land form, and several individual tuff beds can be traced throughout the basin. Tuff near the lake margins consists almost entirely of unaltered rhyolitic glass, and in the center of the lake they are replaced by potassium feldspar, locally associated with searlesite (NaBSi2O5(OH)2). There are scattered occurrences of saline minerals (trona and possibly gaylussite) within the K-feldspar area. Between these areas is the zeolite zone, in which the authigenic minerals are phillipsite-Na, clinoptilolite-Na, erionite-Na, and analcime, in order of abundance.
Similar deposits containing clinoptilolite in the western U.S. include the Pliocene Big Sandy Formation in Mohave County, Arizona, (Sheppard and Gude, 1973); the Miocene Barstow Formation, exposed in the Mud Hills, San Bernardino County, California (Sheppard and Gude 1969); zeolitic tuff in a lacustrine facies of the Gila Conglomerate (Pliocene?) near Buckhorn, Grant County, New Mexico, (Gude and Sheppard 1988); the Pliocene to Holocene lacustrine, Bowie zeolite deposit Cochise and Graham Counties, Arizona (Sheppard et al. 1978); and the huge lake deposits in the Jurassic Morrison Formation, eastern Colorado Plateau (Turner and Fishman 1991).
Clinoptilolite, erionite, and chabazite replace vitric clasts in rhyolitic pyroclastic beds of the Baucarit Formation of Sonora, Mexico (Cochemé et al. 1996, Munch et al. 1996). Farther south in the Sierra Madre Occidental late Miocene volcanism resulted in the accumulation of large deposits of pyroclastic rocks. Some were deposited in closed basins and shallow marine environments, and were diagenetically altered to zeolites, including clinoptilolite (de Pablo-Galán et al. (1996).
Many closed basins, formed by extensional tectonics during the Tertiary in Anatolia, Turkey, had lakes that were saline and alkaline. Contemporaneous volcanic activity contributed pyroclastic and volcaniclastic material to the lake deposits, which in some include lignite or borate beds. Gündogdu et al. (1996) describe the geology and mineralogy of lacustrine deposits containing borate deposits in western Turkey. Beds in the Bigadic, Emet, and Kirka basins include thick sequences of tuff that have been replaced by clinoptilolite-K, which locally has been replaced by analcime, potassium feldspar, and quartz. Analcime and clinoptilolite comprise up to 80% of the mineral matter in coal from the Cayirhan mine, Beypazari, western Turkey (Whateley et al. 1996). Much of this zeolite also occurs as replacement of vitric tuff, caused by reaction with saline and alkaline lake water. In most cases clinoptilolite and analcime replace the vitric component of tuff with relationships similar to those deposits in the western U.S.
Tertiary lacustrine deposits are widespread in Serbia (Obradovic et al. 1995, Obradovic 1988, Gottardi and Obradovic 1978). Tectonic activity developed closed basins, which became the sites of shallow lakes. The lacustrine sediment is largely tuff and volcaniclastic sand and silt accompanied with marl. The volcanic debris, which ranges from rhyodacite to andesite, is replaced by zeolites, commonly clinoptilolite and analcime. In the Slanci basin near Belgrade authigenic minerals are zoned horizontally zoned from fresh glass to montmorillonite to clinoptilolite to analcime (Obradovic 1988). At the base of the section is unaltered sediment from when the early lake was filled with fresh water, and later with a change in climate and lake depth it became sufficiently saline and alkaline to alter the vitric tuff.
Soil and surficial deposits. Clinoptilolite is the most common zeolite occurring in soils, especially those developed from zeolite-bearing parent materials (Ming and Boettinger 2001). Reported occurrences are in the U.S., eastern Europe, Lebanon, and Japan.
Clinoptilolite is widespread in calcareous soils developed on tuffaceous sediment. For example, it is widespread in soils formed on the Catahoula Formation of south Texas (Ming and Dixon 1986) and on volcanic-rich parent units in west Texas (Jacob and Allen 1990). Other occurrences in the U.S. and elsewhere, including Hungry, Bulgaria, Romania, Lebanon, India, and Japan are briefly described by Ming and Boettinger (2001).
Brown et al. (1969) found clinoptilolite or heulandite in soils developed from calcareous, fine-grained siliceous rocks of the Upper Greensand Formation of England. The zeolite may be a diagenetic product from biogenic opal, and was inherited by the soil. A similar occurrence of clinoptilolite was described in calcareous soils of Sicily (Bellanca et al. 1980) and of heulandite in soils developed from the byrozoan chalk in northeastern Denmark (Nørnberg 1990).
Hydrologically open systems. Replacement of rhyolitic pyroclastic rocks in hydrologically open systems generally occurs where pyroclastic deposits have accumulated several hundred meters in thickness and tens of kilometers in lateral extent. Alteration occurs by meteoric water percolating through the deposit, first hydrating glass shards, and deeper into the pile commonly at the water table dissolve glass shards and crystallizing zeolite, mostly clinoptilolite. The authigenic minerals tend to be zoned vertically with slightly altered glass and smectite upper most, clinoptilolite and smectite in a middle layer, and commonly potassium feldspar and analcime with smectite/illite at the base. A common consequence of the hydrologically open system type of replacement is the change in bulk composition by leaching of upper levels and precipitation in lower levels. By contrast replacement processes in hydrologically closed systems produce lateral variations in authigenic mineralogy accompanied by little bulk composition change.
The first well described deposit of the hydrologically open type is the John Day Formation of central Oregon, USA. (Hay 1963). The Oligocene John Day Formation consists of 700 to 1000 m of rhyolitic and dacitic tuff and welded tuff, exposed over an area of at least 5000 km2. In the thicker parts of the formation the upper 500 m contains glass, smectite, and opal, but no zeolite. Clinoptilolite, smectite, and celadonite are abundant in the lower 500 m. Other authigenic minerals include erionite, mordenite, heulandite, and K-feldspar. The contact between the upper and lower parts tends to be planar and cuts across stratification. Hay (1963) describes this boundary as being a cavernous zone 2 to 20 mm thick with many of the vitric particles having been dissolved away. The replacement of glass shards apparently occurs by dissolution of the glass and followed by precipitation of clinoptilolite. K/Ar dates on authigenic K-feldspar deep in the section indicates that most authigenic reactions had occurred by early Miocene, before deposition of the top of the section had been completed.
Another well investigated example of the alteration of a thick pile of pyroclastic rocks is Yucca Mountain along the southwest edge of the Nevada Test Site in Nye County, Nevada, USA. Here a thick sequence of Miocene pyroclastic flow tuff, and lesser amounts of flow breccia, lava, and tuff, has accumulated in a basin at least 1800 m thick in the center and thinning to about 1200 m to the east. These products were erupted from the nearby Timber Mountain-Oasis Valley caldera complex to the north. Because a large block of these rocks is being considered as a repository for radioactive waste, the mountain has been extensively drilled, and the cores studied in detail. Broxton et al. (1987) divide the diagenetic alteration into four mineral zones. Zone 1 is the upper most 375-384 m that consists of unaltered glass with minor smectite, opal and scattered occurrences of heulandite-Ca and clinoptilolite-Ca. In Zone 2, the next 700 to 480 m in depth, the main alteration mineral is clinoptilolite with minor mordenite and lesser amounts of opal, K-feldspar, quartz, and smectite. Zone 3, 400 to 98 m thick, consists of analcime, K-feldspar, and minor calcite and smectite. Zone 4, at the deepest levels, typically has authigenic albite, K-feldspar, and quartz. All zones thin toward the north, the direction of the caldera complex.
The White River sequence, described by Lander and Hay (1993), is an extensive formation largely a non-lacustrine accumulation of rhyolitic tuffs and reworked ash and exposed in Wyoming, South Dakota, and Nebraska, USA. The glass has altered to smectite, opal-CT, and clinoptilolite. Smectite is abundant throughout the formation, but the distribution of opal-CT and clinoptilolite is variable.
Petrova et al. (1987) describe zeolite occurrences at Tushleg and Tsagan-Tsab in southern Mongolia. The upper Jurassic and lower Cretaceous sections consist of 85 to 110 m of tuffaceous sandstone, siltstone, peat-bearing siltstone, and conglomerate with fresh water fossils (fish, insects, and ostracods). Other large deposits occur in Siberia.
Flood and Taylor (1991) and Flood (1995) describe two provinces in eastern Australia in which Carboniferous marine volcaniclastic deposits contain abundant clinoptilolite. The New England province (Early Carboniferous) consists largely of non-marine silicic ignimbrite and tuff units interbedded with fluvial and glacio-fluvial conglomerate, sandstone, and mudstone. Most of the zeolite-rich units are in water-lain and ash-fall tuff. The Late Devonian and Early Carboniferous Drummond basin to the north has a similar stratigraphy, and the zeolite occurs in similar units. Such units contain up to 60 weight percent zeolite, which is in both basins a heulandite group mineral. Because of the fine-grained nature of the zeolite crystals, microprobe analyses have large charge-balance errors, but show that it is either heulandite-Ca or clinoptilolite-Ca. It is, nonetheless, remarkable that the zeolite has remained unchanged since diagenesis.
Other important deposits of zeolitized ignimbrite sheets occur in southwestern Cuba and south Mexico (Popov et al. 2006).
Deep marine sediment. Phillipsite and analcime were discovered by Murray and Renard (1891) in deep-sea sediment collected by the Challenger Expedition. Following several discoveries in the 1960’s, a great many more occurrences were found in core from holes drilled by the Deep Sea Drilling Project. Clinoptilolite-K with phillipsite-K and only a few other zeolites (analcime, harmotome, erionite, and laumontite) are among the authigenic silicate minerals in many different kinds of deep-sea sediment. In general sediments to which the name deep-sea is applied are very fine-grained, occurring in ocean basins well distant from continents and having accumulated at very low rates, 5 m per 106 years.
A great proportion of the clinoptilolite occurrences are in pelagic clay and siliceous microfossil ooze, and chalk. Much less occur in volcaniclastic sediment. In general these sediments are very fine-grained, consisting of various clay minerals, calcite, and poorly crystallized silica minerals, such as opal-CT. Clinoptilolite crystals are euhedral and 2 to 40 µm in length, and comprise 10 to 20% of the sediment. Clinoptilolite has been found in core from all parts of all ocean basins (see Boles, 1977, Iijima 1978, and Kastner and Stonecipher 1978). In general phillipsite occurs in the shallowest levels from very near the water-sediment interface to several hundred meters depth, and clinoptilolite occurs in the deeper and older sections. The zeolite occurrences in deep-sea sediment recovered from legs 1 through about 40 have been reviewed by Stonecipher (1976), Iijima (1978), Boles (1977), and Kastner and Stonecipher (1978).
Even though crystals of clinoptilolite separated from deep-sea sediment are euhedral and clean, their small size (5 to 15 μm) makes obtaining high quality analyses difficult with the electron probe microanalyzer. Good data have been published by Stonecipher (1978) and Boles and Wise (1978). Almost all analyzed samples are clinoptilolite-K, but a few are clinoptilolite-Ca and clinoptilolite-Na still with abundant K, showing that this zeolite has preferentially fractionated K from the interstitial water. Abundant calcite with the nanofossil chalks may account for those few samples with dominant Ca.
There is no compelling textural evidence demonstrating a precursor phase for deep-sea clinoptilolite, such as volcanic glass. A few examples have been found of clinoptilolite replacing siliceous radiolarian tests (Fan and Zemmels 1972). Boles and Wise (1978) show that the conversion of phillipsite to clinoptilolite requires only the addition of water and silica. Therefore, they suggest that phillipsite forms initially in deep-sea sediment as a silica-deficient, metastable phase. With time it dissolves in the pore fluids, and the more stable phase, clinoptilolite, crystallizes where sufficient silica is available.
Diagenesis of mafic lava flows.
The clinoptilolite series minerals are much less common than heulandite in basaltic rocks, but significant occurrences are known. Clinoptilolite-K (the type clinoptilolite) occurs as colorless crystals in altered amygdaloidal basalt breccia exposed on a ridge trending northeast of Hoodoo Peak, Park County Wyoming, USA. (Pirsson 1890). Clinoptilolite-Na forms blocky, crystals in fractures in deeply weathered porphyritic olivine basalt flow-breccia along the north shore of Kamloops Lake, British Columbia, Canada, and in fractures of basalt near Agoura, Santa Monica Mountains, Los Angeles County, California, USA. (Wise et al. 1969). At both localities it is associated with ferrierite and chalcedony. Clinoptilolite-Ca is also known from a few rare occurrences in basalt, such as Kuruma Pass, Fukushima Prefecture, Japan (Koyama and Takéuchi 1977) and in the brecciated tholeiitic basalt of the Columbia River Group, near Altoona, western Wahkiakum County, Washington, USA., where it forms color, blocky crystals in small vesicles and breccia cavities and is associated with mordenite, dachiardite-Na, ferrierite-Na, siderite, and smectite (Wise and Tschernich 1976). Most of these occurrences are in slightly altered host rocks, showing no evidence of pervasive hydrothermal alteration.
Hydrothermal alteration
Active or recently geothermal systems. Replacement of rhyolite tuff intercalated with siltstone and diatomite in lake sediment of the Taupo Volcanic Zone, New Zealand, has produced significant deposits of mordenite and clinioptilolite (Brathwaite 2003). The zeolite deposits are associated with sinter, hydrothermal eruption breccias and silicified fault breccias that represent surface or near-surface geothermal activity. Mordenite and clinoptilolite occur in the lower-T (60–110°C) parts of some active or recently active geothermal systems elsewhere in the Taupo Volcanic Zone.
Uses:
The occurrence of clinoptilolite as fine-grained crystals in massive beds allows it to be mined and used in applications where large amounts of material are required. Significant deposits of clinoptilolite occur in several countries, especially in Bulgaria, Hungary, Cuba, Mexico, Russia, Italy, Jordan, and the United States.
Many applications exploit the ion exchange properties, which are attractive for agronomy, horticulture and soil remediation where the zeolite can be added to chemical fertilizers to improve the soil chemical and physical properties for plant growth, to increase fertilizer efficiency and to reduce the leaching of nutrients. Treatment of effluents containing radioactive contaminants or other heavy metals Other applications make used of the molecular sieve properties, to trap or separate gases in agriculture (e.g. ammonia).
Surfactant (surface active agent) modification of clinoptilolite enhances its use as a sorbent for various solutes not or little affected by the untreated zeolite (Bowman 2003). Some examples are CrO42-, organic molecules (e.g. benzene), and certain biological pathogens.
References:
Akizuki, M., Kudoh, Y., Nakamura, S. 1999. Growth texture and symmetry of heulandite-Ca from Poona, India. Can. Mineral. 37, 1307-1312.
Alberti, A. 1972. On the crystal structure of the zeolite heulandite. Tschermaks Mineral. Petrogr. Mitt. 18, 29-146.
Alberti, A. 1975. The crystal structure of two clinoptilolites. Tschermaks Mineral. Petrogr. Mitt. 22, 25-37.
Alberti, A. and Vezzalini, G. 1983. The thermal behavior of heulandites: A structural study of the dehydration of the Nadap heulandite. Tschermaks Mineral. Petrogr. Mitt. 31, 259-270.
Aleksiev, B and Djourova, E.G. (1975) On the origin of zeolitic rocks. C. R. Acad. Bulg. Sci. 28, 517-520.
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