International Journal of Earth Sciences

, Volume 103, Issue 8, pp 2273–2300 | Cite as

Hydrothermal alteration and zeolitization of the Fohberg phonolite, Kaiserstuhl Volcanic Complex, Germany

  • Tobias Björn Weisenberger
  • Simon Spürgin
  • Yann Lahaye
Original Paper

Abstract

The subvolcanic Fohberg phonolite (Kaiserstuhl Volcanic Complex, Germany) is an economic zeolite deposit, formed by hydrothermal alteration of primary magmatic minerals. It is mined due to the high (>40 wt%) zeolite content, which accounts for the remarkable zeolitic physicochemical properties of the ground rock. New mineralogical and geochemical studies are carried out (a) to evaluate the manifestation of hydrothermal alteration, and (b) to constrain the physical and chemical properties of the fluids, which promoted hydrothermal replacement. The alkaline intrusion is characterized by the primary mineralogy: feldspathoid minerals, K-feldspar, aegirine–augite, wollastonite, and andradite. The rare-earth elements-phase götzenite is formed during the late-stage magmatic crystallization. Fluid-induced re-equilibration of feldspathoid minerals and wollastonite caused breakdown to a set of secondary phases. Feldspathoid minerals are totally replaced by various zeolite species, calcite, and barite. Wollastonite breakdown results in the formation of various zeolites, calcite, pectolite, sepiolite, and quartz. Zeolites are formed during subsolidus hydrothermal alteration (<150 °C) under alkaline conditions. A sequence of Ca–Na-dominated zeolite species (gonnardite, thomsonite, mesolite) is followed by natrolite. The sequence reflects an increase in \(\log [(a_{{{\text{Na}}^{ + } }} )/(a_{{{\text{H}}^{ + } }} )]\) and decrease in \(\log [(a_{{{\text{Ca}}^{2 + } }} )/(a_{{{\text{H}}^{ + } }}^{2} )]\) of the precipitating fluid. Low radiogenic 87Sr/86Sr values indicate a local origin of the elements necessary for secondary mineral formation from primary igneous phases. In addition, fractures cut the intrusive body, which contain zeolites, followed by calcite and a variety of other silicates, carbonates, and sulfates as younger generations. Stable isotope analysis of late-fracture calcite indicates very late circulation of meteoric fluids and mobilization of organic matter from surrounding sedimentary units.

Keywords

Zeolite Götzenite Pectolite Sepiolite Hydrothermal alteration Kaiserstuhl Volcanic Complex 

Introduction

Hydrothermally altered phonolitic rocks are of economic interest in the Kaiserstuhl Volcanic Complex (KVC), SW Germany (Fig. 1), due to the occurrence of zeolites, dominantly natrolite-group minerals and analcime. The pronounced cation exchange and molecular sieve capabilities of zeolites and zeolite-rich rocks make them an important economic target for mining (Pabalan and Bertetti 1999, 2001). These properties have widespread industrial application in water softening (Kallo 2001), catalysis (Weitkamp 2000), remediation of soils and soil quality (Leggo and Ledesert 2001; Leggo et al. 2010; Mercurio et al. 2010, 2012), wastewater treatment (Kallo 2001; Leggo and Ledesert 2001; Leggo et al. 2010), as additive in the cement industry (Kassautzki 1983; Colella et al. 2001; Hauri 2006), or as agent in solar energy heating and energy storage (Tchernev 2001).
Fig. 1

a Simplified geologic map of the Kaiserstuhl Volcanic Complex (after Wimmenauer 1963). The phonolite-hosted Fohberg zeolite deposit is marked. b Outline of Germany and Baden-Württemberg (gray) showing the location of the KVC

Zeolites are tectosilicates characterized by an open three-dimensional framework of (Si, Al)O4 tetrahedra. The tetrahedra form a network of open channels containing molecular water and charge-balancing cations of alkali and alkaline earth metals. Their distinctive crystal structures result in the ability to hydrate/dehydrate reversibly and to exchange cations with aqueous solutions. Zeolites are among the most common products of chemical interaction between groundwater and the Earth’s crust during diagenesis (Boles and Coombs 1977; Boles and Franks 1979; Boles and Surdam 1979), hydrothermal mineral alteration (Kristmannsdóttir and Tómasson 1978; Neuhoff et al. 2000; Weisenberger and Selbekk 2009), and low-grade metamorphism (Weisenberger and Bucher 2010, 2011; Bucher and Weisenberger 2013). They form in low-temperature (<250 °C), low-pressure (<200 MPa), water-saturated, or fluid-rich environments. The copious amounts of silica, aluminum, and alkali and alkaline earth cations necessary for the formation of most zeolites are commonly derived from the dissolution of volcanic glass and from the alteration of feldspar and feldspathoid minerals, which react with an aqueous fluid with low CO2 tolerance at high pH conditions (Zen 1961; Weisenberger and Bucher 2010, 2011).

The most diverse zeolite occurrences are found in large igneous provinces, like in the Deccan Province in India (Sukheswala et al. 1974), or in the North Atlantic Igneous Province (Neuhoff et al. 2000). Nevertheless, large uniform zeolite occurrences of economic interest are associated with altered pyroclastic rocks/volcanic tuffs (Ibrahim and Hall 1996; Ibrahim 2004; Machiels et al. 2014), volcanoclastic sandstones (Boles and Surdam 1979; Surdam and Boles 1979), and pervasively altered lavas and intrusive bodies (Ferguson and Edgar 1978, Fuentes et al. 2004; Weisenberger and Spürgin 2009).

Sodic zeolite species like natrolite-group minerals (natrolite, mesolite, gonnardite, thomsonite) and analcime are often associated with peralkaline to alkaline volcanic products. In particular, natrolite has been known for a long time, mainly as “fibrous zeolite.” Klaproth (1803) introduced the name “natrolite” for a mineral from Hohentwiel, Hegau volcanic field, Germany. At this classic locality, natrolite occurs in a highly fractured, fine-grained, phonolitic neck, filling a fracture network with aggregates showing a characteristic white–red zonation pattern (Vetter 1938). In addition, the host rocks are highly altered and zeolitized (Vetter 1938).

Sequential precipitation of gonnardite and natrolite in mafic rocks of Kahrizak, Iran, and Disko–Nuussuaq, West Greenland, is associated with a change in fluid composition. The \(a_{{{\text{Na}}^{ + } }}/ a_{{{\text{H}}^{ + } }}\) ratio increases with time during diagenesis and very low-grade metamorphism. The change in fluid composition forms the paragenesis chabazite, thomsonite, gonnardite, and natrolite, whereas the fluid composition itself is buffered by the low Si and low Ca lavas (Neuhoff et al. 2006; Kousehlar et al. 2012).

Natrolite-group minerals are also formed as late-stage products of deuteric alteration of alkaline to peralkaline intrusions, like nepheline phonolites, nepheline syenites, and syenitc pegmatite dikes (Deer et al. 2004; Andersen et al. 2010; Schilling et al. 2011), where primary magmatic minerals react with water-rich solutions that separate from the same body of magma at a late stage during cooling. In the Mont Saint–Hilaire peralkaline complex (Quebec, Canada), Schilling et al. (2011) observed analcime and natrolite in syenitic rocks. These sodium-bearing zeolites were formed in a post-magmatic hydrothermal stage, associated with a continuous decrease in \(a_{{{\text{SiO}}_{{ 2 , {\text{aq}}}} }}\) and increase in \(a_{{{\text{H}}_{ 2} {\text{O}}}}\) by the consumption of primary igneous nepheline (Schilling et al. 2011). Stable isotopic data of late-stage carbonates indicate two fluid sources and mixing of them; a magmatic component and a carbon and oxygen isotopic signature indicating the surrounding limestone as fluid source (Schilling et al. 2011).

Several phonolite intrusion bodies were emplaced in the KVC (Fig. 1). Nevertheless, the degree and style of alteration differs between the intrusive bodies, which has a direct effect on the economic potential of the specific phonolites. For instance, the Kircherg phonolite (Fig. 1) on the western flanks of the KVC, near Niederrotweil, is only slightly altered and contains fresh sodalite-group minerals in addition to secondary calcite and sodic zeolites (natrolite, analcime) in the groundmass and in fissures (Wimmenauer 1962). Phonolites in the central subvolcanic part of the KVC are generally altered and show various degrees of zeolitization (Wimmenauer 1962). In contrast, the Fohberg phonolite near Bötzingen (Figs. 1, 2) is mined due to intense alteration and the formation of secondary phases, mainly zeolites. The excellent qualities of the Fohberg phonolite, used in concrete industry and many other applications, are directly related to the high degree of alteration and zeolite content (Kassautzki 1983; Hauri 2006). Other phonolitic bodies in the eastern KVC (Fig. 1) show comparable alteration patterns and may be of potential economic interest. Therefore, the description and interpretation of the alteration processes in the Fohberg phonolite is needed to establish a genetic model as a premise for predictions at other localities in the KVC or elsewhere.
Fig. 2

a Fohberg phonolite intrusion. b View from the center of the quarry into north direction. The phonolite intrusion shows a discordant contact to overlying ocher Pleistocene loess sediments. People for scale. c Enlargement of rectangle given in b. Essexitic dike, approximately 2 m in width is crosscutting the phonolite intrusion and indicates magmatic activity after phonolite emplacement. d Alteration is most obvious along white natrolite (Ntr) veins and open mineralized fractures. The grayish rock matrix is also highly affected by zeolitization. Hammer for scale

In this study, we reconstruct the hydrothermal evolution of the Fohberg phonolite, eastern KVC, as example of an economic zeolite deposit. We present new geochemical, mineralogical, and textural data of the primary igneous mineral assemblage as well as of deuteric and low-temperature alteration phases. Further, we provide a genetic model for the hydrothermal replacement of primary magmatic phases by secondary minerals, and information about fluid and element sources.

Geological setting

The KVC is located in the central-southern segment of the Upper Rhine Graben, southwestern Germany (Fig. 1), which is part of the European Cenozoic Rift System. The Miocene KVC is the only larger volcanic edifice in the Upper Rhine Graben. Alkaline and carbonatitic rocks erupted along deep-rooting faults in a disrupted crustal segment (Bourgeois et al. 2007) at the intersection of two prominent fault zones (Hüttner 1996; Schreiner 1996).

Numerous highly silica-undersaturated alkaline dikes and diatremes of smaller volume than the KVC occur locally in the Upper Rhine Graben (e.g., Mahlberg melilitite) and along the graben shoulders, i.e., in the crystalline complexes of the Black Forest, Vosges, and Odins Forest, as well as in adjacent Mesozoic–Early Cenozoic sedimentary blocks. They are a part of the Cenozoic volcanic province in Central Europe, which formed in the course of the Alpine continent–continent collision (cf. Wedepohl et al. 1994). These graben volcanic and subvolcanic rocks of olivine–melilititic and olivine–nephelinitic composition are the unfractionated products of low-percentage partial mantle melts (Keller 2001). Magmatism in the area of the present Upper Rhine Graben can be traced back to Campanian ages (81 Ma, Lippolt et al. 1974).

The Kaiserstuhl Volcanic Complex (KVC)

The KVC covers an area of 16 × 12 km northwest of Freiburg and rises up to 270 m compared to the surrounding alluvial plain, forming a distinct morphological high. A north–south-oriented pre-volcanic horst structure in the eastern Kaiserstuhl, mainly comprising Paleogene marls, sandstones, and limestones (Wimmenauer 2003), is partly overlain by effusive and explosive volcanic rocks, and is penetrated by subvolcanic intrusions (Fig. 1). Extrusive rocks of the KVC were erupted from various volcanic centers and formed a complex strato-cone or a volcanic field (Keller 2001). Parasitic vents are found in the southwest (Münsterberg/Breisach) and northwest (Limberg/Sasbach) of the KVC (Fig. 1).

Pursuant to their outcrop and petrographic characteristics, the volcanic rocks belong to different groups: (1) lava flows, generally characterized by a porphyric texture of augite phenocrysts with a primary vitreous matrix, predominantly tephrites. (2) Pyroclastic rocks, e.g., ash-/lapillituffs, agglomerates, mainly of tephritic composition, whereas minor phonolitic pyroclastics are found and a rarely carbonatites (Henkenberg, Kirchberg) (Keller 1981). (3) Porphyric dikes with essexitic, phonolitic, or carbonatitic affinity, also intermediate compositions, and highly evolved rocks, most prominent in the inner KVC and radially oriented toward the center. (4) Subvolcanic intrusive rocks of larger dimensions, fine- to medium-grained, especially essexite, phonolite, and carbonatite. (5) Intrusion-related diatreme breccias of subvolcanic niveau.

The magmatic rocks of the alkaline rock–carbonatite complex are supposed to be derived from two different parental magmas (Keller 1984, 2001; Schleicher et al. 1991). A primary olivine–nephelinite magma exposed at the Limberg–Lützelberg complex near Sasbach (Fig. 1) resembles those found at several places in the Upper Rhine Graben. This olivine–nephelinite magma generates slightly fractionated basanites (limburgite), melilite-bearing rocks (bergalite), and carbonatites. The second primary magma is a hypothetical, initially fractionated K-basanite, which shows signs of crustal contamination. It is not exposed in the KVC. Fractionation of this second source material led to two petrographically distinctive rock “clans.” The essexitic clan (Wimmenauer 1957, 1959a; Kim 1985) comprises leucite– and olivine–tephrites, and phonolitic tephrites, which form the major part of the exposed KVC. Continued fractionation caused the formation of rocks belonging to the phonolitic clan (Wimmenauer 1962), e.g., phonolites, syenites, and evolved leucocratic dikes. The alternating deposition of tephritic and phonolitic tuff beds indicates simultaneous activity of these magma systems (Wimmenauer 1962).

Volcanic activity in the KVC started in the Middle Miocene ranging from 19.0 to 15.3 Ma. Activity started with the eruption of olivine nephelinites (19.0 ± 1.6 Ma, Baranyi et al. 1976), followed by the deposition of dominantly tephritic rocks (18.2–16.5 Ma) and the emplacement of various subvolcanic intrusions and dikes during the postulated volcanic main phase between 18.4 and 15.3 Ma (compiled in Wimmenauer 2003). Some of the youngest surface volcanics forming the parasitic Limberg–Lützelberg complex were erupted at 16.2 ± 0.2 Ma (Kraml et al. 2006).

Since the early days of its geological and mineralogical exploration, the KVC is well known for its wealth in zeolites. A detailed description of the zeolite distribution in relation to the host rock is given in Weisenberger and Spürgin (2009). Nevertheless, phonolite intrusions in the eastern KVC, namely the Fohberg phonolite, are the only local zeolite occurrences of economic interest today. Therein, natrolite is the predominant zeolite-group mineral.

Analytical methods

Mineral compositions were determined at the University of Oulu using an electron microprobe JEOL JXA-8200. Operating conditions were 15 kV acceleration voltage and 15 nA beam current with counting times of 10 s. Zeolites were analyzed with a defocused beam (20 μm). Other minerals were analyzed with a beam diameter of 5 or 10 μm. Na and K were measured first, to minimize the effect of Na and K loss during determination. Since zeolites lose water when heated, the crystals were mounted in epoxy resin to minimize the loss of water. Natural and synthetic standards were used for calibration. The charge balance of zeolite formulas is a reliable measure for the quality of the analysis. It correlates with the extent of thermal decomposition of zeolites during microprobe analysis. A useful test is based on the charge balance between the non-framework cations and the amount of tetrahedral Al (Passaglia 1970). Analyses are considered acceptable if the sum E% = 100 × [Al − (Na + K) − 2(Mg + Ca + Sr + Ba)]/[(Na + K) + 2(Mg + Ca + Sr + Ba)] of the charge of the extra-framework cations (Ca2+, Sr2+, Na+, and K+) is within 7 % of the framework charge.

In situ Sr isotope analyses were performed by laser ablation MC–ICP–MS using a Nu Plasma HR multi-collector inductively coupled plasma mass spectrometer and a Photon Machine Analyte G2 laser microprobe at the Geological Survey of Finland in Espoo. Samples were ablated in He gas (gas flows = 0.4 and 0.1 l/min) within a HelEx ablation cell (Müller et al. 2009). All analyses were made in static ablation mode using the following parameters: beam diameter: 110 μm, pulse frequency: 10 Hz, and beam energy density: 2.07 J/cm2. The MC–ICP–MS was equipped with 9 Faraday detectors and amplifiers with 1011 Ω resistors. During the laser ablation, the data were collected in static mode (84Sr–Kr, 85Rb, 86Sr–Kr, 87Rb–Sr, 88Sr). Measured isotope ratios were corrected for instrument fractionation using an exponential law and a 86Sr/88Sr value of 0.1194. The isobaric interference of 87Rb on 87Sr was monitored and corrected using the 85Rb ion signal and a value of 0.38571 for the 87Rb/85Rb ratio. The isobaric interference of 86Kr on 86Sr was corrected using a 30 s background measurement, preceding every ablation. The accuracy of the laser ablation protocol was verified throughout the day of measurement by repeated analysis of an in-house plagioclase standard from a megacryst of the Cameroon volcanic chain (sample Mir a, Rankenburg et al. 2004). The laser ablation parameters were similar on the samples and the standards. During the course of this study, the average 87Sr/86Sr value obtained was 0.70308 ± 0.00007 (2σ, n = 5), similar to the TIMS value of 0.70311 ± 0.0001 (2σ, Rankenburg et al. 2004).

The composition of stable carbon and oxygen isotopes in carbonates was obtained at the Leibniz Laboratory for Radiometric Dating and Stable Isotope Research, Kiel University, using a Thermo Finnigan Kiel IV preparation device interfaced with a Thermo Finnigan MAT 253 mass spectrometer. The measured carbon dioxide is prepared from the carbonates by reaction with phosphoric acid (99 % H3PO4) for 4 min at 74 °C. The external precision of oxygen and carbon isotope data from three laboratory-internal standards (calibrated against NBS 19, and NBS20) is <0.1 ‰ and <0.06 ‰, respectively. In comparison with the laboratory-internal standards, the measured carbonates from this study are less homogenous as indicated by the elevated differences from their replicates.

Field relations

The approximately 600 × 450 m large Fohberg phonolite in the southeastern Kaiserstuhl (Figs. 1, 2) is a shallow subvolcanic stock, which intruded in Oligocene marls of the Pechelbronn formation (Wimmenauer 2003). The upper contact of the phonolite intrusion to the overlying Pleistocene loess unit is discordant (Fig. 2). The phonolite stock is cut by a dike of porphyritic, black essexite (Fig. 2b, c).

A large variety of xenoliths is found within the Fohberg phonolite. Xenoliths can be as big as 10 cm in diameter, but smaller xenoliths are more common. Xenolith compositions range from a variety of primitive igneous cumulates, crystalline basement rocks to the whole suite of Mesozoic and Cenozoic sedimentary rocks and often show a metasomatic reaction rim at the contact to the phonolitic host rock.

Petrography

The Fohberg phonolite is a holocrystalline, fine-grained subvolcanic rock with a weak porphyritic texture (Fig. 3). It is gray in color (Fig. 2b–d) but it also appears in grayish-green and grayish-brown varieties. The porphyritic texture is characterized by wollastonite and aegirine–augite phenocrysts, which can be up to 2–4 mm in length. Feldspathoid phenocrysts are of smaller grain size, <1 mm in diameter, and completely replaced by pseudomorphic zeolites and calcite. Ti-bearing andradite (varietal name “melanite”) occurs as minor phenocryst phase.
Fig. 3

Photomicrographs showing characteristic primary orthomagmatic and alteration textures. a Euhedral aegirine–augite laths adjacent to andradite in a matrix of K-feldspar and calcite. b Aegirine–augite encases strongly corroded andradite. Aegirine–augite shows color zonation from light green in cores to dark green rims. Fibrous wollastonite phenocrysts showing characteristic shape, with arrowhead-like terminations. Matrix consists of K-feldspar and zeolites. c Fine- to medium-grained euhedral to subhedral aegirine–augite crystals showing skeletal texture with a strong embayment of their grain boundaries. The texture indicates the formation of aegirine–augite by the consumption of andradite. d Andradite + aegirine–augite aggregate surrounding euhedral feldspathoid mineral pseudomorphic replaced by calcite, adjacent to a primary wollastonite lath, replaced by natrolite and calcite. Matrix consists of K-feldspar and natrolite. e Phenocrysts of former feldspathoid minerals replaced by zeolites characterized by hexagonal, rhombic dodecahedral, and rectangular cross sections. Wollastonite is replaced by zeolite and/or calcite preferentially along its apex. f Natrolite replacing feldspathoid minerals. The euhedral shape is highlighted by submicroscopic inclusions that are enriched along former grain boundaries. g Natrolite replacing feldspathoid mineral showing enrichment of brownish inclusions of unknown composition, surrounded by K-feldspar laths and zeolites as matrix replacement. In contrast to the feldspathoid mineral replacement, which appears cloudy, matrix zeolites are transparent. h Aggregate of fine natrolite needles replacing feldspathoid minerals, surrounded by larger natrolite crystals and K-feldspar. Mineral abbreviations: Adr andradite, Cal calcite, Kfs K-feldspar, Ntr natrolite, Wo wollastonite, Agt aegirine–augite. Mineral abbreviations in all figures according to Whitney and Evans (2010) and Bucher and Grapes (2011)

The equigranular groundmass is fine- to very fine-grained and mainly composed of K-feldspar. In addition, zeolites and calcite occur in different proportion as alteration products in the groundmass. As accessory phases following minerals occur in various proportions, either as primary or secondary phases: götzenite, apatite, barite, sepiolite, pectolite, and titanite.

Fracture petrography

The Fohberg phonolite intrusion is cut by a fracture network (Fig. 2d). Extensional joints may be a result of subsolidus cooling and shrinkage or of fluid-induced fracturing as a result of boiling and/or overpressuring during post-magmatic hydrothermal activity. The fracture network consists of several prominent sets of subvertical and, much less distinctive, sets of flat dipping joints. According to Stober (1955), the joint sets are mainly related to cooling, although some may reflect regional structural directions. The macroscopic fracture network is highly regular over the outcrop size with a fracture aperture in the centimeter scale and fracture spacing of some decimeter to several meter scales. Individual fractures can be followed over several tens of meters.

The fractures are healed partially or totally with a secondary mineral assemblage dominated by zeolites (natrolite, gonnardite, thomsonite) and calcite. A series of other infrequent phases may be found, e.g., apophyllite, fluorite, hyalite, strontianite, celestine, barite, clay minerals, and hydrocarbons (Wimmenauer 1959b; Marzi 1983). In cases where relative mineral successions are evident, zeolites always appear as first and calcite as last (probably excluding clays and hydrocarbons) precipitate.

Mineral compositions

K-feldspar

K-feldspar occurs as fine-grained (<150 μm) euhedral to subhedral, interstitial microlites between different phenocrysts (Fig. 3g). These microlites appear to surround phenocryst phases (feldspathoid minerals and aegirine–augite) in a trachytic texture. K-feldspar microlites often depict Carlsbad twinning. In some places, K-feldspar seems to be mechanically fractured and generated fracture porosity is filled with secondary zeolites (Fig. 4). Rarely zeolites replace K-feldspar (Fig. 4). The composition of K-feldspar in the phonolitic rock is homogenous with An0.0–0.5Ab11–19Or89–81 (Table 1; Fig. 5a).
Fig. 4

Backscattered electron images of characteristic alteration textures of wollastonite. a Wollastonite replaced by calcite and natrolite. b Gonnardite replacing wollastonite. K-feldspar mechanically broken due to hydraulic fracturing. c Calcite and quartz replacing wollastonite. d Wollastonite breakdown to sepiolite and calcite. Matrix consists of calcite, which formed as replacement of feldspathoid minerals and contains small barite inclusions. A larger barite grain is also visible. e Wollastonite breakdown to sepiolite and calcite adjacent to aegirine–augite. Calcite as replacement product contains quartz inclusions. f Natrolite replacing wollastonite. Feldspathoid minerals are replaced by gonnardite. K-feldspar appears mechanically broken and corroded. Götzenite together with K-feldspar is clustered around gonnardite. Mineral abbreviations: Adr andradite, Brt barite, Cal calcite, Kfs K-feldspar, Ntr natrolite, Qz quartz, Sep sepiolite, Ttn titanite, Wo wollastonite, Agt aegirine–augite, Go götzenite, Gon gonnardite

Table 1

Representative compositions of aegirine–augite (core and rim), andradite, K-feldspar, and titanite from Fohberg

Sample

Aegirine–augite

FP2

FP2

FP2

FP2

FP2

FP2

NBK

NBK

NBK

NBK

SP12 Cc06

SP12 Cc06

Analysis no

2

3

9

16

26

27

41

42

55

56

5

6

 

Core

Rim

  

Core

Rim

Core

Rim

Core

Rim

Core

Rim

SiO2

50.29

49.43

49.88

49.91

49.98

49.63

49.19

48.84

48.67

48.98

48.96

49.52

TiO2

0.27

0.58

0.97

0.37

0.18

0.55

0.38

0.39

0.35

0.54

0.44

0.73

Al2O3

2.25

0.77

0.46

1.33

2.10

0.53

2.55

1.05

2.65

0.61

2.63

0.51

Fe2O3

4.96

13.82

20.10

7.33

6.67

15.97

4.83

8.09

5.50

15.10

6.81

18.03

FeO

9.25

12.57

7.36

9.78

7.94

11.04

10.13

10.96

8.47

10.48

8.28

9.70

MnO

0.96

1.36

1.41

1.25

0.99

1.53

1.13

1.30

0.93

1.45

0.94

1.40

MgO

8.56

0.88

0.26

6.61

8.81

0.39

7.40

5.39

7.96

1.18

7.98

0.27

CaO

22.27

13.29

10.03

20.78

21.80

11.90

21.33

19.64

22.64

13.83

22.33

10.84

Na2O

1.39

5.85

8.27

2.23

1.57

6.75

1.58

2.50

1.26

5.89

1.49

7.43

K2O

0.00

0.00

0.02

0.01

0.03

0.01

0.01

0.01

0.01

0.00

0.00

0.00

ZrO2

0.15

0.12

0.02

0.03

0.01

0.19

0.00

0.04

0.00

0.03

0.06

0.17

SrO

0.06

0.05

0.00

0.25

0.00

0.14

0.00

0.13

0.00

0.00

0.00

0.09

Totala

100.41

98.72

98.81

99.89

100.14

98.66

98.57

98.38

98.48

98.08

99.93

98.69

 

Formula based on 4 cations and 8 oxygen

Si

1.92

1.98

1.98

1.94

1.91

1.99

1.92

1.94

1.90

1.97

1.88

1.98

Al

0.10

0.04

0.02

0.06

0.09

0.03

0.12

0.05

0.12

0.03

0.12

0.02

Ti

0.01

0.02

0.03

0.01

0.01

0.02

0.01

0.01

0.01

0.02

0.01

0.02

Fe3+

0.14

0.42

0.60

0.21

0.19

0.48

0.14

0.24

0.16

0.46

0.20

0.54

Mg

0.49

0.05

0.02

0.38

0.50

0.02

0.43

0.32

0.46

0.07

0.46

0.02

Fe2+

0.30

0.42

0.24

0.32

0.25

0.37

0.33

0.36

0.28

0.35

0.27

0.32

Mn

0.03

0.05

0.05

0.04

0.03

0.05

0.04

0.04

0.03

0.05

0.03

0.05

Ca

0.91

0.57

0.43

0.86

0.89

0.51

0.89

0.84

0.95

0.60

0.92

0.46

Na

0.10

0.45

0.64

0.17

0.12

0.52

0.12

0.19

0.10

0.46

0.11

0.58

K

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

Zr

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

Sr

0.00

0.00

0.00

0.01

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

Sample

Andradite

K-feldspar

Titanite

FP2

NBK

NBK

SP12 Cc06

FP2

FP2

NBK

SP12 Cc06

SP12 Cc06

SP12 Cc06

Analysis no

4

11

53

4

1

15

10

7

10

11

SiO2

31.53

31.35

31.24

32.36

64.51

64.13

64.15

63.61

29.81

30.12

TiO2

7.73

8.78

7.91

6.70

0.02

0.02

0.00

0.01

36.15

37.04

Al2O3

0.94

0.82

0.62

0.76

18.09

18.57

17.84

18.09

0.16

0.17

Fe2O3

24.06

21.53

22.39

25.30

      

FeO

1.40

2.73

1.61

1.29

0.71

0.60

0.76

0.52

2.36

1.90

MnO

0.76

0.80

0.81

0.75

0.01

0.00

0.00

0.00

0.03

0.04

MgO

0.24

0.24

0.24

0.24

0.00

0.00

0.00

0.01

0.01

0.01

CaO

32.04

31.67

31.68

32.13

0.01

0.00

0.01

0.00

27.32

27.95

Na2O

0.25

0.23

0.25

0.28

1.37

1.25

1.40

1.64

0.30

0.36

K2O

0.01

0.00

0.00

0.00

14.95

14.86

14.60

14.31

0.00

0.00

ZrO2

0.29

0.41

0.25

0.33

    

0.20

0.10

SrO

0.00

0.00

0.00

0.00

0.19

0.46

0.12

0.27

0.56

0.42

Totala

99.42

98.76

96.99

100.13

99.92

100.01

98.91

98.50

97.25

98.29

 

Formula based on 16 cations and 24 oxygen

Formula based on 8 oxygen

Formula based on 5 oxygen

Si

5.38

5.39

5.45

5.48

2.99

2.97

3.00

2.98

1.02

1.01

Al

0.19

0.17

0.13

0.15

0.99

1.01

0.98

1.00

0.01

0.01

Ti

0.99

1.14

1.04

0.85

0.00

0.00

0.00

0.00

0.93

0.94

Fe3+

3.09

2.78

2.94

3.22

      

Mg

0.06

0.06

0.06

0.06

0.00

0.00

0.00

0.00

0.00

0.00

Fe2+

0.20

0.39

0.24

0.18

0.03

0.02

0.03

0.02

0.07

0.05

Mn

0.11

0.12

0.12

0.11

0.00

0.00

0.00

0.00

0.00

0.00

Ca

5.86

5.83

5.92

5.83

0.00

0.00

0.00

0.00

1.00

1.01

Na

0.08

0.08

0.09

0.09

0.12

0.11

0.13

0.15

0.02

0.02

K

0.00

0.00

0.00

0.00

0.88

0.88

0.87

0.86

0.00

0.00

Zr

0.02

0.03

0.02

0.03

    

0.00

0.00

Sr

0.00

0.00

0.00

0.00

0.01

0.01

0.00

0.01

0.01

0.01

aTotals include traces of Ce, La, Nb, Zr

Aegirine–augite

Aegirine–augite occurs as euhedral to subhedral elongated laths and anhedral grains with a distinct zonation pattern (Fig. 3). The cores show pale-green color under plane-polarized light, whereas the rims are strongly pleochroic from dark green, to brownish-green, and greenish yellow (Fig. 3c, g). Very often, the fine- to medium-grained euhedral to subhedral aegirine–augite crystals exhibit a skeletal texture with a strong embayment of their grain boundaries (Fig. 3). Aegirine–augite is often associated with andradite and titanite (Fig. 6b) and encases strongly corroded andradite grains. The chemical composition of aegirine–augite and its variation is presented in Fig. 5 and Table 1. In terms of endmember proportions, aegirine–augite is Di16–50Hd25–44Agt09–59 and Jd1–6Agt4–68Quad29–92. The jadeite component ranges between 1 and 6 mol% (Fig. 5c). The optical zoning pattern is reflected by a sharp chemical zoning separating core from rim (Fig. 3b, e, g). It is caused by an increase in aegirine component, reflected by an increase in Fe3+, Fe2+, and Na+ mol%, and a decrease in Ca and Mg mol% from core to rim, respectively (Fig. 5, Table 1). Titanium (<1.5 wt% TiO2) and manganese occur as minor elements in aegirine–augite. Both show a gently increasing pattern from core to rim (Table 1). Aegirine–augite cores exhibit an average Ti and Mn content of 0.01 and 0.03 mol%, whereas corresponding rims show an average Ti and Mn content of 0.02 and 0.05 mol%, respectively.
Fig. 5

Ternary plots showing primary igneous compositions: a K-feldspar compositions fall along the K-feldspar–albite solid solution series, whereas an anorthite component is absent. Isotherms are taken from Fuhrman and Lindsley (1988). b Ternary plot of the Y-site contents of garnets. Fohberg garnet is classified as andradite according to Grew et al. (2013). c, d Ternary diagrams classifying clinopyroxene of the Fohberg phonolite. c Composition of clinopyroxene cores falls into the field of Quad and aegirine–augite. Later formed clinopyroxene has a higher aegirine component. d Ternary plot showing clinopyroxene compositions in the system diopside–hedenbergite–aegirine. Dashed lines indicate the trends of data from other localities of alkaline rocks (Ilimaussaq, Marks and Markl 2001; North Qôroq, Coulson 2003; Alnö, Vuorinen et al. 2005)

Andradite

Ti-rich andradite (“melanite”) forms fine-grained (0.1–1 mm in diameter), greenish brown and dark brown, subhedral to anhedral micro-phenocrysts (Fig. 3a–c). It shows a strong sign of alteration characterized by a cauliflower-like or skeletal shape (Fig. 3a–c). Andradite is associated with aegirine–augite and encased by it, suggesting andradite is replaced by clinopyroxene (Figs. 3a–c, 6b). Andradite is characterized by a high Ti content ranging from 0.84 to 1.14 mol% (6.52–8.78 wt% TiO2, Fig. 5; Table 1). The chemical composition is homogenous (Table 1) with an average composition of Ca5.84Mg0.07Fe0.242+Mn0.11Na0.08Fe3.013+Ti0.97Al0.17Si5.45O24. Zirconium occurs in traces from 0.02 to 0.03 mol% (0.25–0.41 wt% ZrO) (Table 1).
Fig. 6

Backscattered electron images of characteristic alteration textures. a Zeolites (natrolite and gonnardite) replacing feldspathoid minerals. In contrast to thin section microphotographs, no euhedral grain shape is visible in BSE. Cores of aegirine–augite show euhedral shape, whereas the rim is subhedral. b Andradite is generally surrounded by aegirine–augite, which contains inclusions of andradite and titanite. c Enlargement of rectangle in a. Zeolites replacing feldspathoid minerals. Porous gonnardite is overgrown by natrolite. Small (<3 μm) inclusions of barite occur in zeolites. K-feldspar shows porosity. d Andradite + aegirine–augite aggregate surrounding calcite replacing euhedral feldspathoid mineral. A primary wollastonite lath is replaced by natrolite and calcite. Matrix consists of K-feldspar and natrolite. Black area left of the wollastonite lath represents porosity gained during replacement process. e, f Götzenite associated with K-feldspar is clustered around gonnardite replacing feldspathoid minerals. K-feldspar appears mechanically broken. Mineral abbreviations: Adr andradite, Brt barite, Cal calcite, Kfs K-feldspar, Ntr natrolite, Ttn Titanite, Wo wollastonite, Agt aegirine–augite, Go götzenite, Gon gonnardite

Wollastonite and replacement products

Wollastonite is a major phase (up to 12 vol%, Albrecht 1981) in the Fohberg phonolite, occurring as euhedral, fibrous phenocrysts up to 1 cm in length with both ends forming an arrowhead-like shape (Figs. 3b–c, 4, 7). Along the tip of the arrowheads, wollastonite shows alteration to various secondary phases (Figs. 3b–c, 4, 7). The chemistry of wollastonite is close to endmember composition (Table 2). Alteration can proceed to total replacement of wollastonite by secondary phases. Such replacement is very heterogeneous, and several secondary mineral assemblages are observed. (1) Pectolite (Ca2NaH(SiO3)3) replaces wollastonite (Fig. 7), the average composition of pectolite being close to endmember composition (Ca2Fe0.03Mn0.04Na0.93K0,01H(Si1.01O3)3, Albrecht 1981). (2) Wollastonite breaks down forming the assemblage calcite plus quartz (Fig. 4c). (3) Breakdown of wollastonite can be accompanied by the formation of the zeolite species gonnardite or natrolite (Fig. 4a, b, f). Additional calcite occurs, but may be absent in cases where relictic wollastonite is present (Fig. 4). (4) Figure 4d, e shows the breakdown of wollastonite to a phyllosilicate-like phase and calcite, adjacent to aegirine–augite. This phase has low analytical totals (Table 2) and a composition suggestive for sepiolite (Mg4Si6O15(OH)2·6H2O). The average composition of sepiolite is Mg4.02Ca0.16Mn0.07Fe0.06Na0.01K0.01Si5.49Al0.46O15(OH)2·6H2O.
Fig. 7

Photomicrographs showing characteristic primary orthomagmatic and alteration textures. a Euhedral wollastonite partly replaced by pectolite. b Götzenite, filling interstitial space between K-feldspar laths and feldspathoid pseudomorphs. Mineral abbreviations: Pct pectolite, Wo wollastonite, Go götzenite

Table 2

Representative compositions of wollastonite, sepiolite, and calcite from Fohberg

Sample

Wollastonite

Sepiolite

Calcite

NBK

NBK

NBK

NBK

NBK

NBK

NBK

NBK

NBK

NBK

NBK

FP2

FP2

FP2

NBK

SP12 Cc06

Analysis no

1

2

9

22

23

30

31

16

32

33

34

20

23

44

62

2

           

Rock

Rock

Rock

Fracture

Fracture

SiO2

50.87

50.49

50.27

50.75

50.25

50.74

51.50

49.59

47.63

47.98

47.64

0.00

0.00

0.00

0.00

0.00

TiO2

0.01

0.03

0.02

0.02

0.00

0.01

0.02

0.00

0.01

0.01

0.00

0.00

0.01

0.00

0.00

0.00

Al2O3

0.00

0.00

0.03

0.00

0.00

0.01

0.02

3.76

2.89

3.51

3.54

0.00

0.00

0.02

0.00

0.00

FeO

0.96

0.90

0.75

0.73

0.69

0.75

0.75

0.33

0.50

1.20

0.44

0.42

0.04

0.11

0.00

0.02

MnO

0.63

0.62

0.87

0.79

0.89

0.43

0.41

1.70

0.53

0.19

0.51

1.27

0.68

0.81

0.03

0.01

MgO

0.18

0.18

0.11

0.12

0.14

0.19

0.18

24.49

23.51

22.84

23.94

0.09

0.01

0.02

0.03

0.00

CaO

46.79

46.76

46.65

46.67

46.51

46.98

47.13

1.66

1.34

1.13

1.00

55.15

55.07

56.14

55.01

55.37

Na2O

0.01

0.02

0.00

0.01

0.01

0.03

0.04

0.07

0.05

0.00

0.05

0.04

0.10

0.00

0.00

0.01

K2O

0.00

0.01

0.01

0.00

0.01

0.01

0.00

0.12

0.09

0.10

0.06

0.01

0.02

0.00

0.00

0.00

ZrO2

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.01

0.09

0.00

0.00

0.00

0.33

0.00

0.00

0.00

SrO

0.00

0.00

0.10

0.00

0.13

0.05

0.00

0.00

0.01

0.00

0.00

     

Cl

       

0.03

0.09

0.09

0.10

     

CO2a

           

44.49

44.02

44.66

43.22

43.49

Totalb

99.49

99.12

98.85

99.20

98.65

99.26

100.15

81.80

76.82

77.07

77.40

101.51

100.47

101.75

98.29

98.92

 

Formula based on 3 oxygen

Anhydrous formula unit composition based on 11 oxygen

Formula based on 1 cation

Si

0.99

0.99

0.99

0.99

0.99

0.99

1.00

5.43

5.52

5.54

5.47

0.00

0.00

0.00

0.00

0.00

Al

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.49

0.40

0.48

0.48

0.00

0.00

0.00

0.00

0.00

Ti

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

Mg

0.01

0.01

0.00

0.00

0.00

0.01

0.01

4.00

4.06

3.93

4.10

0.00

0.00

0.00

0.00

0.00

Fe

0.02

0.01

0.01

0.01

0.01

0.01

0.01

0.03

0.05

0.12

0.04

0.01

0.00

0.00

0.00

0.00

Mn

0.01

0.01

0.01

0.01

0.01

0.01

0.01

0.16

0.05

0.02

0.05

0.02

0.01

0.01

0.00

0.00

Ca

0.98

0.98

0.99

0.98

0.98

0.99

0.98

0.19

0.17

0.14

0.12

0.97

0.98

0.99

1.00

1.00

Na

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.01

0.01

0.00

0.01

0.00

0.00

0.00

0.00

0.00

K

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.02

0.01

0.01

0.01

0.00

0.00

0.00

0.00

0.00

Zr

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.01

0.00

0.00

     

Sr

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

Cl

       

0.01

0.02

0.02

0.02

     

aCalculated

bTotals include traces of Ce, La, Nb, Zr

Feldspathoid minerals and replacement products

Feldspathoid minerals occur as phenocrysts ranging between 0.1 and 1 mm in diameter (Fig. 3). It is the most abundant phenocryst phase in the Fohberg phonolite. These phenocrysts are totally altered to a pseudomorphic aggregate of fibrous spherulitic zeolites (natrolite and gonnardite) and carbonate, leaving pseudomorphs with dominant hexagonal, but also rhombic dodecahedral and rectangular cross sections (Figs. 3, 4). Zeolite species are the dominant secondary replacement products of feldspathoid minerals and are present constantly with 45 volume percent (Albrecht 1981). In contrast, the calcite content is variable from totally absent to modal volume proportions of >2 % (Albrecht 1981). Zeolites replacing feldspathoid minerals often contain small (<3 μm) inclusions of barite (Fig. 6c). In rare cases, larger irregular barite is found (Fig. 4d). Pseudomorphs are very conspicuous due to the enrichment of very fine-grained, submicroscopic brownish inclusions along the primary euhedral crystal faces, which are again overgrown by pseudomorphic replacement products (zeolites and calcite) (Fig. 4f, g). Direct identification of the inclusions is not possible due to the submicroscopic nature. In backscattered electron images (Figs. 4, 6), no inclusions are visible and no pseudomorphic shape can be recognized, due to the fact that zeolites and calcite occur also as matrix replacement product.

Zeolites

Zeolites are the major secondary replacement product in the Fohberg phonolite, replacing feldspathoid minerals, wollastonite, and unspecified matrix phases. Additional, zeolites occur in fractures (Fig. 8), lining fracture walls as first phase of a zeolite–calcite-dominated assemblage. Fracture zeolites form long, thin needles elongated along the c-axis, terminated by (111) faces, as well as radial aggregates and compact masses. Color varies from colorless to white, and light yellowish to light red–brown. Individual crystals can have a size of more than 10 mm in length, but smaller needles (<1 mm) are more common. Zeolites as replacement products in the phonolite fabric occur as fine-grained mineral aggregates. Host rock zeolites as replacement products (Figs. 4, 6) as well as fracture zeolites (Fig. 8) show a change in mineralogy from a Ca–Na zeolite species to natrolite (Na16Al16Si24O80·16H2O). Figure 6c shows porous Ca–Na zeolite followed by natrolite. In fracture samples, Ca–Na zeolite, which is cloudy under the optical microscope and appears to be in a broken texture, is followed by natrolite that contains small (<10 μm) inclusions of calcite (Fig. 9). X-ray diffraction analysis suggests that the Ca–Na zeolite in fractures is thomsonite (Ca8Na4Al20Si20O80·24H2O), gonnardite (Na,Ca)12–16(Al,Si)40O80.24H2O), or both, depending on the sample. Chemical analyses of the Ca–Na zeolite from one of these fractures presented in Figs. 10, 11 and Table 3 are close to gonnardite compositions reported elsewhere (e.g., Deer et al. 2004). Furthermore, the Ca–Na zeolite mesolite (Na16Ca16Al48Si72O240·64H2O) occurs together with natrolite predominantly along sites where natrolite is porous (Fig. 9a). Mesolite grows parallel to the natrolite c-axis, forming laths with a width below 2 μm and a length of maximum 40 μm (Fig. 9a). It is not distinguishable if this pattern is the result of epitactical growth, natrolite replacement, or an exsolution texture. Chemical analyses of zeolites were obtained on natrolite and gonnardite from the phonolite alteration assemblages, and from the same phases occurring as fracture-filling (Figs. 3, 4, 6, 8; Table 3). For both, natrolite and gonnardite, the composition of phonolite alteration phases shows a wider spread than the data for the corresponding fracture phases (Figs. 10, 11). Fracture natrolite shows very little compositional variation. Sodium is the dominant extra-framework cation, with a minor amount of Ca, in average 0.04 atoms per formula unit (apfu, natrolite formula normalized to 80 oxygens), and traces of Fe, Mg, Mn, Sr, and K. The average TSi (Si/(Si + Al)) ratio is 0.600 with a range from 0.597 to 0.601(Table 3). Natrolite occurring as replacement product in phonolite shows a wider compositional variation. Beside the dominant extra-framework cation Na, Ca occurs in average of 0.26 apfu with a range of 0.03–0.62 apfu. Other elements, including Fe, Mg, Mn, Sr, and K, do not significantly get incorporated. The average TSi ratio is 0.604 within a range from 0.592 to 0.610 (Table 3). It cannot be excluded that the wider compositional variation of natrolite as replacement product is an artifact of measuring, due to submicroscopic inclusions of calcite and/or mesolite analyzed with a defocused electron beam. Gonnardite compositions of replacement assemblages are characterized by Na and Ca as major extra-framework cations (Figs. 10, 11; Table 3), with an average Na content of 9.05 apfu (gonnardite formula normalized to 80 oxygens) in the range between 8.26 and 10.37 apfu and an average Ca content of 3.96 apfu in the range between 3.58 and 4.32 apfu. Gonnardite contains Mg, on average 0.12 apfu; however, values up to 0.60 apfu Mg are observed. Another significant minor element is Sr, with an average of 0.32 apfu, ranging from 0.20 to 0.52 apfu; Fe, Mn, and K only occur in traces. The TSi ratio of gonnardite as replacement product is lower than TSi ratio values for natrolite, with an average of 0.548, ranging between 0.515 and 0.573 (Table 3). Gonnardite as replacement product shows large chemical variations and deviates from the ideal gonnardite series. In contrast, fracture gonnardite is close to the gonnardite series proposed by Ross et al. (1992) (Fig. 11). Beside Na and Ca, only Sr occurs in significant values in fracture gonnardite (0.33 apfu), whereas Mg and Fe are not found. The TSi ratio of fracture gonnardite is 0.55 (Table 3).
Fig. 8

Photomicrographs showing characteristic fracture paragenesis. a Fibrous zeolite aggregates followed by euhedral calcite growing on the apex of the aggregate. Host rock is located below the bottom edge of the picture, and the fracture is subparallel to the bottom edge. Early formed zeolite species is gonnardite, which is epitaxial overgrown by natrolite. b Multiple generations of calcite as fracture-filling. Host rock is exposed at the lower part of the image, dominated by black color. (c) View parallel to the c-axis of natrolite, showing rectangular cross section of natrolite surrounded by calcite. Natrolite contains parallel intergrowths as described in Akizuki and Harada (1988). Mineral abbreviations: Cal calcite, Ntr natrolite, Gon gonnardite

Fig. 9

Backscattered electron images of characteristic alteration textures. a Natrolite replacing feldspathoid minerals. Mesolite grows or replaces natrolite parallel to the c-axis. Natrolite contains microscopic calcite inclusions. b Zeolite-filled fracture. Early gonnardite is overgrown by natrolite. Calcite inclusions occur in natrolite, but are absent in gonnardite. Mineral abbreviations: Cal calcite, Mes mesolite, Ntr natrolite, Gon gonnardite

Fig. 10

Geochemical characterization of zeolites from Fohberg; shown are compositions of fracture-filling and host rock natrolite and gonnardite. a Extra-framework cation distribution. b Si-R2+-R+ plot. Compositional fields (dashed lines) for natrolite-group minerals and thomsonite taken from Deer et al. (2004)

Fig. 11

Composition of natrolite and gonnardite from Fohberg for fracture and host rock samples expressed in terms of Ca and Al per 20 framework oxygens. The increase in Al content of gonnardite from the rock matrix is accompanied by a Na increase. The dashed lines represent the gonnardite series and thomsonite solid solution (ss) proposed by Ross et al. (1992)

Table 3

Representative compositions of zeolites from Fohberg

Sample

Natrolite

Sp12 Apo2

Sp12 Apo2

Sp12 Apo3

Sp12 Apo3

SP12 Cc06

NBK

NBK

FP2

FP2

NBK

NBK

SP12 Cc06

SP12 Cc06

Analysis no

c_7

d_1

c_1

c_3

3

63

64

33

34

36

38

8

9

Fracture

Fracture

Fracture

Fracture

Fracture

Fracture

Fracture

Rock

Rock

Rock

Rock

Rock

Rock

SiO2

46.04

47.35

46.64

46.64

47.35

46.35

46.39

47.11

47.98

47.71

46.87

47.64

47.07

Al2O3

26.38

26.66

26.52

26.35

26.69

26.10

26.13

26.23

26.26

26.02

26.06

26.45

26.10

FeO

0.00

0.02

0.00

0.01

0.08

0.00

0.01

0.01

0.01

0.00

0.03

0.08

0.06

MgO

0.00

0.00

0.02

0.00

0.00

0.01

0.01

0.00

0.01

0.15

0.00

0.00

0.00

CaO

0.07

0.38

0.18

0.01

0.03

0.00

0.09

0.06

0.13

0.55

0.82

0.12

0.07

SrO

    

0.00

0.00

0.02

0.00

0.01

0.00

0.08

0.00

0.00

Na2O

15.61

15.51

15.54

15.61

16.08

15.38

15.21

15.70

15.72

14.83

14.14

15.30

15.41

K2O

0.00

0.00

0.00

0.02

0.02

0.02

0.02

0.04

0.03

0.02

0.02

0.02

0.02

TiO2

0.00

0.00

0.00

0.00

0.00

0.02

0.01

0.00

0.00

0.01

0.00

0.00

0.00

MnO

    

0.03

0.00

0.01

0.00

0.02

0.00

0.01

0.00

0.00

Totals

88.09

89.93

88.89

88.64

90.27

87.93

87.91

89.21

90.20

89.38

88.06

89.61

88.72

 

Anhydrous formula unit composition based on 80 oxygen

Si

23.93

24.08

24.00

24.07

24.04

24.11

24.11

24.17

24.32

24.37

24.27

24.26

24.24

Al

16.16

15.98

16.08

16.03

15.97

16.00

16.01

15.86

15.69

15.66

15.91

15.87

15.84

Fe

0.00

0.01

0.00

0.01

0.03

0.00

0.01

0.00

0.01

0.00

0.01

0.04

0.03

Mg

0.00

0.00

0.01

0.00

0.00

0.01

0.01

0.00

0.01

0.12

0.00

0.00

0.00

Ca

0.04

0.21

0.10

0.00

0.02

0.00

0.05

0.03

0.07

0.30

0.45

0.06

0.04

Sr

0.00

0.00

0.00

0.00

0.00

0.00

0.01

0.00

0.00

0.00

0.03

0.00

0.00

Na

15.73

15.30

15.50

15.62

15.82

15.51

15.33

15.62

15.45

14.68

14.19

15.11

15.39

K

0.00

0.00

0.00

0.01

0.01

0.01

0.01

0.03

0.02

0.01

0.01

0.01

0.01

Ti

0.00

0.00

0.00

0.00

0.00

0.02

0.01

0.00

0.00

0.01

0.00

0.00

0.00

Mn

0.00

0.00

0.00

0.00

0.01

0.00

0.00

0.00

0.01

0.00

0.01

0.00

0.00

E %

2.26

1.72

2.30

2.46

0.60

3.00

3.47

1.02

0.42

0.84

4.90

4.10

2.38

Si/(Si + Al)

0.60

0.60

0.60

0.60

0.60

0.60

0.60

0.60

0.61

0.61

0.60

0.60

0.60

Sample

Gonnardite

NBK

NBK

FP2

FP2

NBK

NBK

Analysis no

65

66

11

36

6

59

Fracture

Fracture

Rock

Rock

Rock

Rock

SiO2

41.05

41.31

40.09

38.84

42.52

42.56

Al2O3

28.85

28.43

32.00

30.04

27.93

28.41

FeO

0.00

0.00

0.00

0.08

0.15

0.14

MgO

0.00

0.00

0.03

0.00

0.03

0.32

CaO

7.28

6.26

6.88

7.07

6.83

7.02

SrO

1.11

1.03

0.81

0.81

1.27

1.72

Na2O

8.23

9.13

10.34

9.92

8.40

8.16

K2O

0.02

0.01

0.03

0.01

0.02

0.01

TiO2

0.02

0.01

0.00

0.00

0.00

0.00

MnO

0.03

0.00

0.00

0.05

0.01

0.03

Totals

86.60

86.20

90.35

86.88

87.22

88.37

 

Anhydrous formula unit composition based on 80 oxygen

Si

21.95

22.17

20.73

20.92

22.56

22.34

Al

18.18

17.98

19.51

19.06

17.46

17.57

Fe

0.00

0.00

0.00

0.03

0.07

0.06

Mg

0.00

0.00

0.02

0.00

0.02

0.25

Ca

4.17

3.60

3.81

4.08

3.88

3.95

Sr

0.34

0.32

0.24

0.25

0.39

0.52

Na

8.53

9.50

10.37

10.36

8.64

8.31

K

0.02

0.01

0.02

0.00

0.01

0.01

Ti

0.01

0.01

0.00

0.00

0.00

0.00

Mn

0.02

0.00

0.00

0.02

0.00

0.01

E%

3.41

3.64

5.28

0.19

1.29

−1.04

Si/(Si + Al)

0.55

0.55

0.52

0.52

0.56

0.56

* E% = (100 × [Al − (Na + K) − 2(Mg + Ca + Sr + Ba)]/[(Na + K) + 2(Mg + Ca + Sr + Ba)]

Table 4

Representative compositions of götzenite from Fohberg

Sample

NBK

NBK

NBK

NBK

NBK

NBK

NBK

NBK

NBK

NBK

NBK

NBK

NBK

Analysis no

14

15

16

17

18

43

44

46

45

47

48

49

50

SiO2

31.52

31.00

30.25

30.90

31.27

30.90

30.51

30.50

30.64

30.75

31.45

30.56

31.43

TiO2

8.13

8.11

6.86

8.24

7.90

7.11

7.86

7.89

6.83

7.86

6.92

7.84

7.84

Al2O3

0.02

0.00

0.00

0.00

0.09

0.00

0.02

0.00

0.00

0.00

0.18

0.01

0.05

FeO

0.21

0.15

0.28

0.17

0.22

0.23

0.19

0.26

0.21

0.20

0.26

0.19

0.31

MnO

0.44

0.25

0.18

0.19

0.23

0.28

0.18

0.14

0.16

0.13

0.27

0.21

0.25

MgO

0.00

0.01

0.02

0.01

0.00

0.01

0.00

0.01

0.00

0.00

0.01

0.02

0.00

CaO

37.80

37.13

35.03

36.20

36.74

36.02

35.45

35.29

35.30

35.70

37.19

35.48

36.49

Na2O

5.88

6.00

6.30

5.95

6.22

6.20

6.34

6.23

6.08

6.53

6.21

6.35

6.07

K2O

0.05

0.06

0.03

0.11

0.06

0.06

0.04

0.11

0.04

0.04

0.06

0.02

0.08

ZrO2

1.72

1.06

1.17

0.90

1.57

0.84

0.84

1.23

1.03

0.88

1.49

1.02

1.18

SrO

2.39

2.34

2.55

2.23

2.30

2.41

2.22

2.49

2.32

1.87

2.25

2.56

2.23

La2O3

0.58

1.40

1.85

1.49

0.91

1.72

1.97

1.42

1.68

1.62

1.11

1.54

1.49

Ce2O3

0.89

2.22

3.46

2.51

1.26

3.36

3.46

2.96

3.39

3.51

1.18

3.23

2.20

Nb2O3

     

3.29

2.40

2.96

3.01

2.64

3.21

2.93

3.37

Cl

0.00

0.00

0.00

0.00

0.01

0.00

0.01

0.01

0.00

0.03

0.01

0.00

0.01

F

6.64

6.82

6.60

6.72

6.57

6.80

6.65

6.82

6.79

6.63

6.92

6.76

6.41

−O ≡ F

3.79

3.90

3.77

3.84

3.75

3.89

3.80

3.90

3.88

3.79

3.95

3.87

3.66

Total

93.47

93.70

91.80

92.79

92.58

96.38

95.32

95.43

94.62

95.56

95.79

95.88

96.71

 

Formula based on 8 oxygen

Si

2.00

1.99

2.01

2.00

2.01

1.98

1.97

1.96

2.00

1.97

1.99

1.96

1.98

Al

0.00

0.00

0.00

0.00

0.01

0.00

0.00

0.00

0.00

0.00

0.01

0.00

0.00

Ti

0.39

0.39

0.34

0.40

0.38

0.34

0.38

0.38

0.33

0.38

0.33

0.38

0.37

Mg

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

Fe

0.01

0.01

0.02

0.01

0.01

0.01

0.01

0.01

0.01

0.01

0.01

0.01

0.02

Mn

0.02

0.01

0.01

0.01

0.01

0.02

0.01

0.01

0.01

0.01

0.01

0.01

0.01

Ca

2.57

2.56

2.50

2.52

2.53

2.47

2.45

2.43

2.46

2.45

2.52

2.44

2.46

Na

0.72

0.75

0.81

0.75

0.78

0.77

0.79

0.78

0.77

0.81

0.76

0.79

0.74

K

0.00

0.01

0.00

0.01

0.01

0.00

0.00

0.01

0.00

0.00

0.00

0.00

0.01

Zr

0.05

0.03

0.04

0.03

0.05

0.03

0.03

0.04

0.03

0.03

0.05

0.03

0.04

Sr

0.09

0.09

0.10

0.08

0.09

0.09

0.08

0.09

0.09

0.07

0.08

0.10

0.08

La

0.01

0.03

0.05

0.04

0.02

0.04

0.05

0.03

0.04

0.04

0.03

0.04

0.03

Ce

0.02

0.04

0.07

0.05

0.03

0.07

0.07

0.06

0.07

0.07

0.02

0.07

0.04

Nb

     

0.11

0.08

0.10

0.10

0.09

0.10

0.10

0.11

Cl

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

0.00

F

1.33

1.39

1.39

1.38

1.34

1.38

1.36

1.39

1.40

1.34

1.39

1.37

1.27

Calcite

Calcite occurs similar to natrolite and gonnardite as replacement product in the phonolite and as fracture-filling. As replacement product, the modal content of calcite is highly variable, replacing wollastonite, feldspathoid minerals, and matrix phases (Figs. 4, 6). In zeolite–calcite-filled fractures, calcite always follows natrolite as younger generation (Fig. 8). However, some fractures contain exclusively calcite, which can show all sizes from very fine-grained to crystals several cm in size. The chemical composition of calcite replacing primary phases is characterized by the incorporation of significant contents of MnO (up to 1.27 wt%), FeO (up to 0.42 wt%), and SrO (up to 0.33 wt%) (Table 2). In contrast, fracture calcite shows only traces of these oxides (Table 2).

Accessory minerals

Titanite, apatite, barite, and götzenite occur as accessory minerals, either as primary igneous phases or as products of deuteric alteration. Apatite forms cloudy euhedral crystals with prismatic and hexagonal cross sections. They occur as inclusions in replacement products of feldspathoid minerals, wollastonite, and aegirine–augite. Titanite occurs either as small (<5 μm) euhedral inclusions, or as anhedral grains in association with andradite and aegirine–augite. Barite forms inclusions in feldspathoid replacement products. Common are tiny grains <3 μm (Fig. 6c), but larger aggregates of barite (up to 75 μm in size) have been found (Fig. 4d).

Rare Zr–Ti–Nb disilicates are characteristic accessory minerals of alkaline rocks and carbonatites (Bellezza et al. 2004). Götzenite ((Ca, Zr)4(Ca, Na)4Ca4(Na, Ca)2Ti2(Si2O7)4(F, O, OH)8; Bellezza et al. 2004) from the Fohberg phonolite was first described by Czygan (1973) and is easily distinguishable by its anomalous blue to brownish interference colors (Fig. 7c). Götzenite is the most Ca- and Ti-rich member of the rosenbuschite group (Christiansen et al. 2003). The abundance of götzenite is variable on outcrop scale. Albrecht (1981) observed higher götzenite contents (>1.4 vol%) in higher phonolite levels, whereas in the deeper part of the intrusion götzenite is below 0.5 vol% or even absent. On thin section scale, götzenite appears not to be homogenous distributed. In fact, it occurs in clusters together with K-feldspar around gonnardite pseudomorphoses after feldspathoid phenocrysts (Fig. 6e, f) as well as clusters associated with andradite. The chemical composition of all analyzed götzenite (n = 13) is homogeneous (Table 4). All götzenite analyses systematically have totals below 100 wt%, with fluorine content always lower than 2 apfu (götzenite formula normalized to 8 oxygen and fluorine) (Table 4). This points to the presence of hydroxyl anions substituting for fluorine. In contrast to other roschenbuschite-group minerals like kochite, rosenbuschite, and seidozerite, götzenite does not have separate zirconium polyhedra in its structure (Christiansen et al. 2003), which is reflected by the analyses of the Fohberg samples. Zr is only a minor substituent for Ti in the range of 0.03–0.05 apfu (Table 4). Titanium contents range between 0.33 and 0.40 apfu, Ca ranges between 2.43 and 2.57, and Na varies from 0.72 to 0.81, respectively. Strontium contents range between 0.06 and 0.09.; Al, K, Fe, Mg, and Mn occur in traces (<0.01 apfu). Götzenite contains an average of ~7 wt% rare-earth elements (REE), the most abundant are La2O3, Ce2O3, and Nb2O3 with average values of 1.44 wt% (0.03 apfu), 2.59 wt% (0.05 apfu), and 2.98 wt% (0.09 apfu), respectively, and ranges of 0.58–1.97, 0.89–3.51 and 2.40–3.37 wt%, respectively (Table 4).

Stable isotopes

Stable isotope compositions (δ13CVPDB and δ18OSMOW) were analyzed for 10 samples of calcite fracture-fillings and one sample (SP12-Cc11) of the overlying Pleistocene loess unit (Table 5; Fig. 12). The majority of fracture-filling calcite clusters at δ18OSMOW values of 23.8 ‰ (23.01–24.75 ‰) and δ13CVPDB values of −11.8 ‰ (−13.06 to −10.06 ‰). Samples SP12-Cc2, SP12-Cc3, and SP12-Cc8 differ from this cluster having heavier δ13CVPDB and lighter δ18OSMOW values (Fig. 12). These three samples form a trendline with negative slope pointing to the cluster of the other fracture calcite samples (Fig. 12). The loess sample has an average value of δ13CVPDB −2.8 ‰ and δ18OSMOW 23.5 ‰ (n = 4, Table 5; Fig. 12). The δ18OSMOW value is in the same range as the cluster of fracture calcite, but δ13CVPDB is significantly lighter in the loess sample than in all fracture calcites and also differs from their common trendline.
Table 5

Stable and strontium isotopic compositions of fracture and matrix calcite and fracture natrolite

Sample

Mineral

Cement type

δ18OVPDB (‰)

δ18OSMOW (‰)

δ13CVPDB (‰)

87Sr/86Sr

2σ

Paragenesis

Description

Sp12-Cc01

Cal

Fracture

−7.12

23.57

−12.82

  

Ntr–Cal

Fracture-filling sequence

−7.43

23.25

−13.06

Sp12-Cc02

Cal

Fracture

−9.09

21.54

−7.34

0.704859

0.000045

Ntr–Cal

Aggregate of natrolite followed by euhedral Cal crystals

−8.83

21.8

−7.59

0.704613

0.000019

   

0.704669

0.000038

   

0.704704

0.000047

 

Ntr

Fracture

   

0.703847

0.00031

  

0.704053

0.000138

0.703867

0.000115

 

Cal

Matrix

   

0.704014

0.00005

  

0.704114

0.000024

0.703887

0.000035

 

Ntr

Matrix

   

0.702172

0.000066

  

Sp12-Cc03

Cal

Fracture

−11.45

19.1

−5.65

  

Cal lode

Fine-grained Cal, followed by larger, transparent Cal crystals; large crystals taken for stable isotope analysis

−11.89

18.66

−5.56

Sp12-Cc04

Cal

Fracture

−7.15

23.54

−12.52

  

Cal lode

Layered Cal fracture

−6.96

23.73

−12.40

Sp12-Cc05

Cal

Fracture

−7.16

23.53

−10.71

0.705728

0.000056

Cal lode

Coarse grained Cal fracture fill

−6.32

24.4

−10.61

0.705621

0.000059

   

0.705951

0.000088

   

0.705706

0.000054

   

0.705955

0.000085

   

0.706032

0.000159

   

0.705795

0.000063

 

Cal

Matrix

   

0.704194

0.000015

  

0.704274

0.00003

Sp12-Cc06

Cal

Fracture

−7.34

23.35

−12.05

  

Ntr–Cal

Cal fracture coating

−7.32

23.36

−12.02

Sp12-Cc07

Cal

Fracture

−6.19

24.53

−11.92

  

Cal

Coarse grained, massif Cal fracture fill

−5.98

24.75

−12.08

Sp12-Cc08

Cal

Fracture

−12.68

17.84

−4.97

  

Cal

Layered fibers Cal crystals aligned parallel to fracture wall

−12.91

17.6

−4.72

Sp12-Cc09

Cal

Fracture

−6.49

24.22

−10.06

  

Cal

Coarse grained Cal fracture fill

−6.63

24.08

−10.15

Sp12-Cc10

Cal

Fracture

−7.66

23.01

−12.61

  

Cal

Transparent Cal, comb-like texture

−7.08

23.61

−12.58

Sp12-Cc11

Cal

Fracture

−2.93

23.95

−6.75

  

Loess

Bulk rock sample from overlying Quaternary sedimentary unit

−2.67

23.31

−7.37

−2.86

23.44

−7.25

−2.66

23.24

−7.44

Fig. 12

Stable isotope composition of calcite fracture cement in the Fohberg phonolite. For comparison, various carbonate phases from the KVC are given (including calcite from carbonatites and alkaline silicate rocks, and hydrothermal, secondary, and sedimentary calcites; Hubberten et al. 1988). The rectangle represents the stable isotope composition for carbonatites after Taylor et al. (1967) and Hoefs (1973)

Sr-isotopes

Strontium isotopic ratios (87Sr/86Sr) were analyzed on calcites of the phonolite alteration assemblage and from fracture-fillings of two different samples (Sp12-Cc02, Sp12-Cc05; Table 5; Fig. 13). Additional natrolite fracture-fillings were measured in sample Sp12-Cc02 (Table 5; Fig. 13). Within a single fracture, the 87Sr/86Sr isotopic ratios of calcite and natrolite are homogenous. In sample SP12-Cc02, the average 87Sr/86Sr ratio for fracture calcite is 0.70471 (n = 3; 0.70461–0.70486), whereas in sample SP12-Cc05 the average 87Sr/86Sr ratio for fracture calcite is slightly higher at 0.70583 (n = 7; 0.70571–0.70603) (Table 5). In both samples, corresponding calcite from the phonolite alteration assemblage has slightly lower values than the fracture calcite. Sample SP12-Cc02 has an average 87Sr/86Sr ratio of 0.70401 (n = 3; 0.70389–0.70411) for alteration calcite, whereas sample SP12-Cc05 has an average 87Sr/86Sr ratio of 0.70406 (n = 7; 0.70401–0.70411) (Table 5). Natrolite as fracture phase predating calcite in sample SP12-Cc02 has an average 87Sr/86Sr ratio of 0.70392 (n = 3; 0.70385–0.70405), which is lower than postdating calcite (Table 5). Isotopic ratios of both fracture calcite samples analyzed here are in the same range than other carbonate veins observed in volcanic rocks of the KVC (Wimmenauer 2010).
Fig. 13

87Sr/86Sr ratios for calcite matrix and fracture cements in the Fohberg phonolite (solid symbols). Gray boxes represent primary 87Sr/86Sr ratios for different igneous rocks in the KVC after Schleicher et al. (1990). 87Sr/86Sr ratios for calcite veins and bulk loess samples in the KVC after Wimmenauer (2010) are given as open symbols

Discussion

The conversion of an essentially anhydrous primary mineralogy to a zeolite-dominated and therefore water-rich assemblage (natrolite-group minerals contain about 9.5–14 wt% H2O) on the one hand, and the syn- to post-deformative mineralization and healing of brittle fractures (Fig. 2) on the other hand, point to a hydrothermal regime at subsolidus conditions necessary for the observed zeolitization process. A likely scenario is hydrothermal overprinting of the Fohberg phonolite body during post-magmatic cooling and late-stage circulation of meteoric fluids.

The mineralogy of the Fohberg phonolite can be subdivided to (a) the primary igneous paragenesis, (b) a subsolidus hydrothermal alteration paragenesis in the host rock and fracture system, and c) a low-temperature paragenesis in the fracture system. Metasomatic alteration around xenoliths is limited to the millimeter to centimeter scale and therefore cannot account as major pervasive process in the hydrothermal replacement of the primary phonolite mineralogy.

Primary mineralogical characteristics of the phonolite

Phenocrystic and groundmass minerals form the primary igneous mineral assemblage of the Fohberg phonolite: feldspathoid minerals + K-feldspar + aegirine–augite + wollastonite + andradite + götzenite + apatite ± titanite.

In terms of the observed range of mineral compositions, their evolutionary trends, and mineral relationships, the Fohberg phonolite has similarities with highly evolved alkaline and peralkaline rocks known from several volcanic complexes, like in the East African rift (Zaitsev et al. 2012), the Ilímasussaq complex in Greenland (Markl et al. 2001), or the Mont Saint–Hilaire Complex in Quebec, Canada (Schilling et al. 2011).

Aegirine–augite, andradite, K-feldspar, götzenite, titanite, and apatite are formed during primary igneous crystallization and late-stage re-equilibration. Aegirine–augite is optically and chemically zoned (Fig. 3c, g). We observed two zonation patterns: one with sharp boundaries separating core zones from rims is interpreted to be of primary magmatic origin, whereas a second one showing irregular patches and overgrowths is of late-stage re-equilibration. K-feldspar occurs as late igneous matrix phase with a homogenous composition (Fig. 5). The fact that K-feldspar shows no major sign of breakdown reaction and dissolution suggests that K-feldspar was in equilibrium with altering fluid. Andradite is an early igneous crystallization phase. However, the strong sign of alteration characterized by skeletal and overgrowth textures suggests that andradite broke down during progressing crystallization, forming aegirine–augite and titanite. Late-stage melts enriched in REE and fluorine, and Ti released by andradite breakdown, facilitated the formation of götzenite.

Alteration reactions

Subsolidus alteration in the Fohberg phonolite is marked by the breakdown of feldspathoid minerals and wollastonite. The alteration assemblage is dominated by zeolites (natrolite, gonnardite, thomsonite, mesolite) and calcite. Pectolite, quartz, and barite appear occasionally as replacement products in the phonolite.

It is not evident what feldspathoid mineral (or minerals) occurred as primary igneous phase(s) in the Fohberg phonolite; due to the total replacement by zeolites, no primary remnants have been found. The dominant hexagonal and rhombic dodecahedral cross sections favor a sodalite-group mineral (Fig. 3f–g), whereas the appearance of rectangular cross sections is indicative for nepheline (Fig. 4e–h). On the basis of our observations, it cannot be excluded that both, a sodalite-group mineral and nepheline, are the precursors of zeolites and carbonate. Breakdown of feldspathoid minerals release components necessary for zeolite formation according to the following reactions:
$$2 {\text{NaAlSiO}}_{ 4} ({\text{Nph}}) + {\text{SiO}}_{{ 2,{\text{aq}}}} + 2 {\text{H}}_{ 2} {\text{O}} \Rightarrow {\text{Na}}_{ 2} {\text{Al}}_{ 2} {\text{Si}}_{ 3} {\text{O}}_{ 10} \cdot 2 {\text{H}}_{ 2} {\text{O (Ntr)}}$$
(1)
$${\text{Na}}_{ 8} {\text{Al}}_{ 6} {\text{Si}}_{ 6} {\text{O}}_{ 2 4} {\text{Cl}}_{ 2} ({\text{Sod}}) + 3 {\text{SiO}}_{{ 2 , {\text{aq}}}} + 6 {\text{H}}_{ 2} {\text{O}} \Rightarrow 3 {\text{Na}}_{ 2} {\text{Al}}_{ 2} {\text{Si}}_{ 3} {\text{O}}_{ 10} \cdot 2 {\text{H}}_{ 2} {\text{O (Ntr)}} + 2 {\text{NaCl}}$$
(2)
Reactions (1) and (2) describe the breakdown of the Na-endmember of nepheline and sodalite, respectively, forming natrolite. However, the presence of Ca–Na zeolite species (gonnardite, mesolite, thomsonite) and calcite as secondary phases suggests that the primary feldspathoid minerals were not of pure Na-endmember composition. Reactions (3) and (4) give representative breakdown reactions of Ca-bearing nepheline to gonnardite and natrolite, respectively.
$$\begin{gathered} 2{\text{Na}}_{ 7} {\text{CaAl}}_{ 9} {\text{Si}}_{ 7} {\text{O}}_{ 3 2} ({\text{Nph}}) + 2 {\text{Ca}}^{ 2+ } + 8 {\text{SiO}}_{{ 2,{\text{aq}}}} + 2 4 {\text{H}}_{ 2} {\text{O}} \hfill \\ \quad \Rightarrow {\text{Na}}_{ 10} {\text{Ca}}_{ 4} {\text{Al}}_{ 1 8} {\text{Si}}_{ 2 2} {\text{O}}_{ 80} \cdot 2 4 {\text{H}}_{ 2} {\text{O (Gon)}} + 4 {\text{Na}}^{ + } \hfill \\ \end{gathered}$$
(3)
$$\begin{gathered} 2 {\text{Na}}_{ 7} {\text{CaAl}}_{ 9} {\text{Si}}_{ 7} {\text{O}}_{ 3 2} \, ({\text{Nph)}} + 4 {\text{Na}}^{ + } + 1 3 {\text{SiO}}_{{ 2,{\text{aq}}}} + 1 8 {\text{H}}_{ 2} {\text{O}} \hfill \\ \quad \Rightarrow 9{\text{Na}}_{ 2} {\text{Al}}_{ 2} {\text{Si}}_{ 3} {\text{O}}_{ 10} \cdot 2 {\text{H}}_{ 2} {\text{O (Ntr)}} + 2 {\text{Ca}}^{ 2+ } \hfill \\ \end{gathered}$$
(4)
$$\begin{gathered} 3{{\text{Na}}_{ 6} {\text{Ca}}_{ 2} {\text{Al}}_{ 6} {\text{Si}}_{ 6} {\text{O}}_{ 2 4} \left( {{\text{SO}}_{ 4} } \right)_{ 2} }\left( {\text{Hyn}} \right) + 4 {\text{SiO}}_{{ 2,{\text{aq}}}} + 2 4 {\text{H}}_{ 2} {\text{O}} \hfill \\ \quad \Rightarrow {\text{Na}}_{ 10} {\text{Ca}}_{ 4} {\text{Al}}_{ 1 8} {\text{Si}}_{ 2 2} {\text{O}}_{ 80} \cdot 2 4 {\text{H}}_{ 2} {\text{O (Gon)}} + 8 {\text{Na}}^{ + } + 2 {\text{Ca}}^{ 2+ } + 6 {\text{SO}}_{ 4}^{ 2- } \hfill \\ \end{gathered}$$
(5)
The breakdown of Ca-bearing nepheline leads to the formation of gonnardite under consumption Ca2+ and SiO2,aq and release of Na+ (Reaction 3). Ca2+ and SiO2,aq are provided by the breakdown of wollastonite. The release of Na+ leads to an increase of log[\(a_{{{\text{Na}}^{ + } }}^{{}}\))/(\(a_{{{\text{H}}^{ + } }}^{{}}\))] in the fluid and subsequently to the formation of natrolite (Fig. 14). A similar process is responsible for the formation of mesolite as early fibrous natrolite-group mineral (Fig. 9a). However, the limited amount of sufficient Ca2+ limits the formation of Ca–Na zeolite species. The breakdown of Ca-bearing nepheline without aqueous Ca2+ in access leads to the formation natrolite. Other potential primary feldspathoid minerals are sodalite-group minerals, in particular haüyne (Hyn) with a Ca–Na solid solution. The breakdown of haüyne is given in reaction (5), forming gonnardite and releasing Ca2+, Na+, and SO42− into the fluid. The occurrence of small (<3 μm) barite (BaSO4) inclusions in zeolites (Fig. 6c) suggests that the breakdown of haüyne released SO42− to the fluid.
Fig. 14

Calculated mineral stability diagram at 50 and 100 °C and 10 MPa as a function of cation activity ratios in the Al2O3–Na2O–CaO–SiO2–H2O system, assuming aluminum balance, quartz undersaturation (a(Qz) = 0.95) and a(H2O) = 1. Dashed lines illustrate wollastonite saturation. Natrolite is replaced by analcime at higher aqueous silica activities. Due to the lack of thermodynamic data for gonnardite, mesolite is considered instead. Mineral abbreviations: Cbz chabazite, Grs grossular, Kln kaolinite, Mes mesolite, Ntr natrolite, Thm thomsonite, Wo wollastonite

Calcium is contributed to the fluid during wollastonite breakdown. Wollastonite breakdown results in a variety of secondary alteration assemblages: calcite, calcite–quartz, calcite–natrolite, calcite–gonnardite, pectolite, and sepiolite. The variety of secondary mineral assemblages excludes a fluid with homogenous composition throughout the intrusion. Moreover, it indicates that small-scale variations in fluid composition occur.
$${\text{CaSiO}}_{ 3} \, ({ \text{Wo}}) + {\text{CO}}_{ 2} + {\text{H}}_{ 2} {\text{O}} \Rightarrow {\text{CaCO}}_{ 3} \, ({\text{Cal}}) + {\text{SiO}}_{ 2} \, ({\text{Qz}}) + {\text{H}}_{ 2} {\text{O}}$$
(6)

Reaction (6) is sensitive to the fluid composition in the system CO2–H2O with decreasing temperature (e.g., Bucher and Grapes 2011). Assuming the temperature range of zeolite formation (<200 °C), this results in a fluid with low CO2 tolerance (\(X_{{{\text{CO}}_{ 2} }}\) = <0.01), in accordance with the general low-CO2 formation conditions of zeolite (Zen 1961; Weisenberger and Bucher 2010, 2011). The formation of gonnardite and natrolite during the breakdown of wollastonite depends on local conditions, in particular log[\(a_{{{\text{Na}}^{ + } }}^{{}}\))/(\(a_{{{\text{H}}^{ + } }}^{{}}\))], whereas temperature variations do not play a significant role (Fig. 14).

Pectolite occurs as breakdown product of wollastonite. Pectolite is known to appear in similar geological settings, e.g., the occurrence of pectolite, wollastonite, and götzenite at Mt. Shaheru, north Kivu, Zaire (Sahama and Hytönen 1957), the formation of pectolite during metasomatism in a contact aureole between agpaitic nepheline syenite and sedimentary limestones in the Tamazeght Complex, Morocco (Allah et al. 1998), pectolite formation in a fracture-filling sequence with sepiolite–pectolite–natrolite during alteration of alkaline basaltic rocks at Nuussuaq, West Greenland (Rogers et al. 2006), and as alteration product of REE ore from the Norra Kärr Alkaline Complex, Southern Sweden (Sjöqvist et al. 2013).
$$2 {\text{CaSiO}}_{ 3} \, ({\text{Wo}}) + {\text{SiO}}_{{ 2,{\text{aq}}}} + {\text{H}}_{ 2} {\text{O}} + {\text{Na}}^{ + } \Rightarrow {\text{Ca}}_{ 2} {\text{NaH(SiO}}_{ 3} )_{ 3} \, ({\text{Pct}}) + {\text{H}}^{ + }$$
(7)

Pectolite formation as described in reaction (7) requires Na metasomatism to convert wollastonite to pectolite, Na+ being derived from the breakdown of feldspathoid minerals. The proposed process predicts local pH decrease (via H+ release). Karup-Møller (1969) assumed that pectolite–xonotlite occurring in veins cutting alkaline basaltic rocks in west Greenland has formed under higher temperature conditions at 300 ± 50 °C and pressured less than 30 MPa. The temperature range is in accordance with the observations made by Dutrow et al. (2001) during coupled heat and mass transport calculation, modeling the effect of an alkaline monchiquite dike intrusion into sedimentary rocks, yielding a maximum temperature for pectolite formation between 359 and 250 °C, and Yagi et al. (1968), which synthesized pectolite at 300 °C and <30 MPa.

Sepiolite is a hydrous Mg-rich clay mineral occurring in an assemblage with wollastonite and aegirine–augite in the Fohberg phonolite (Fig. 4d, f). Reaction (8) describes the formation of sepiolite (Table 2) from an aqueous solution. In contrast to all other low-temperature secondary minerals, sepiolite contains Mg. The fact that sepiolite is always associated with aegirine–augite, wollastonite, and calcite (Fig. 4d, f) suggests that Mg2+ is derived from the breakdown of aegirine–augite. Sepiolite and calcite form during wollastonite breakdown according to reaction (9).
$$4 {\text{Mg}}^{ 2+ } + 6 {\text{SiO}}_{{ 2 , {\text{aq}}}} + 1 1 {\text{H}}_{ 2} {\text{O}} \Rightarrow {\text{Mg}}_{ 4} {\text{Si}}_{ 6} {\text{O}}_{ 1 5} ({\text{OH}})_{ 2} \cdot 6 {\text{H}}_{ 2} {\text{O (Sep)}} + 8 {\text{H}}^{ + }$$
(8)
$$6 {\text{CaSiO}}_{ 3} \, ({\text{Wo}}) + 4 {\text{Mg}}^{ 2+ } + 1 1 {\text{H}}_{ 2} {\text{O}} + 6 {\text{CO}}_{ 2} \Rightarrow {\text{Mg}}_{ 4} {\text{Si}}_{ 6} {\text{O}}_{ 1 5} ({\text{OH}})_{ 2} \cdot 6 {\text{H}}_{ 2} {\text{O (Sep)}} + 8 {\text{H}}^{ + } + 6 {\text{CaCO}}_{ 3} \, ({\text{Cal}})$$
(9)

Although sepiolite–palygorskite-group minerals are most widespread and well known in soils, lakes, or shallow seas associated with a Mediterranean to semiarid climate (e.g., Callen 1984; Birsoy 2002), sepiolite–palygorskite-group minerals are also reported in association with hydrothermal alteration fluids in volcanic rocks (e.g., Karup-Møller 1969; Rogers et al. 2006). Sepiolite is formed in environments where the aqueous solution has a low aluminum concentration, high silica concentration, and high pH (Isphording 1973; Birsoy 2002). Laboratory synthesis of sepiolite suggests that sepiolite is unstable at hydrothermal conditions >300 °C (e.g., Güven and Carney 1979). This is in agreement with results by Mumpton and Roy (1958). They proposed that montmorillonite and saponite are more stable than sepiolite or palygorskite at 200 °C. Sepiolite–palygorskite-group minerals associated with pyroxenes and amphiboles are likely to have formed by the alteration of these minerals at low temperatures (Mumpton and Roy 1958).

There is no evidence of a chronological sequence of pectolite, sepiolite, and zeolites, due to the only minor appearance of pectolite and sepiolite and the lack of growth or crosscutting relationships. However, observations from veins sets in basaltic rocks at Nuussuaq, West Greenland, suggest the sequence sepiolite–pectolite–natrolite (Rogers et al. 2006). Whether this sequence is a result of temperature drop and/or changes in fluid chemistry is unknown.

Thermobarometric conditions

A temperature estimate for gonnardite and natrolite formation is hard to assess. In contrast to low-grade meta-basalt alteration, where conditions of formation are usually accomplished by comparison with “calibrated” natural analogues (Neuhoff et al. 1997, 2000; Weisenberger and Selbekk 2009)—e.g., active geothermal systems in Iceland (Kristmannsdóttir and Tómasson 1978), fluid inclusion thermometry (Gilg et al. 2003), Si–Al substitution thermobarometers in zeolites, like stilbite (Fridriksson et al. 2001) and analcime (Neuhoff et al. 2004)—comparative data for assemblages of natrolite-group minerals are rare.

Rogers et al. (2006) estimated temperatures not exceeding 80 °C for similar mineral assemblages in west Greenland, comparing them with the occurrence at County Antrim, Ireland (Walker 1960). The mineral paragenesis analcime and natrolite in western Greenland corresponds to temperatures less than 100 °C (Neuhoff et al. 2006). Therefore, the temperature for zeolite formation in the Fohberg phonolite is reasonably below <100 °C.

The depth of emplacement of the Fohberg phonolite and the resulting lithostatic pressure is unknown. In general, pressure changes are far less effective than temperature changes on zeolite stability. In the following discussion, we use a pressure of 10 MPa, which is a reasonable estimate according to the general geology of the eastern Kaiserstuhl, corresponding to a depth of <1 km for hydrostatic pressure or ~300 m for lithostatic pressure conditions, respectively. The selected pressure regime is below the boiling temperature of H2O for the zeolite stability field. Nevertheless, during earlier cooling stages at higher temperature, boiling may have caused the intense fracturing of the Fohberg phonolite.

Fluid evolution

Fluid evolution can be illustrated in the systems Al2O3–SiO2–Na2O–H2O (Fig. 15) and Al2O3–SiO2–CaO–Na2O–H2O (Fig. 14) in terms of mineral stability as function of aqueous species (Ca2+, Na+, SiO2,aq, H2O) at low temperature and pressure. Mineral stabilities were obtained using the program SUPCRT92 (Johnson et al. 1992), employing the slop98 database (http://geopig.asu.edu/sites/default/files/slop98.dat) and thermodynamic data from Helgeson et al. (1978) and Neuhoff (2000). Reactions were calculated considering low quartz activity and aluminum conservation in the solid phases.
Fig. 15

Quantitative \(a_{{{\text{SiO}}_{ 2} }}\)\(a_{{{\text{H}}_{2} {\text{O}}}}\) diagrams for selected Na–Al silicates at constant pressure (10 MPa) and different temperatures. Gray areas are stability fields for zeolite species. Dashed lines represent the lower limits of quartz saturation. Mineral abbreviations: Ab albite, Anl analcime, Ne nepheline, Ntr natrolite, Qz quartz

Zeolites in the matrix of the Fohberg phonolite are obviously replacing feldspathoid minerals. The limited range of zeolite compositions (Fig. 10), as well as the vast dominance of zeolites in the fissure and alteration assemblages (beside calcite), implies that the decay of primary phases other than feldspathoid minerals plays only a subordinate role in the composition of the percolating fluids. The partial breakdown of wollastonite may account for Ca2+ release to the fluid, whereas the decrease in pH during pectolite formation after wollastonite, as indicated by textural observations, has only an effect on the mineral scale environment, due to fluid buffering. Similarly, the limited decay of aegirine–augite results in the immediate consumption of available Mg2+ during the formation of Mg-rich clay (sepiolite).

Figure 15 illustrates the change in fluid composition during the hydrothermal replacement as a function of H2O and aqueous silica activity, respectively. Phase equilibria between nepheline, albite, natrolite, and analcime allow the estimation of the activities of SiO2 and H2O at fixed pressure and temperature conditions. At a subsolidus temperature of 250 °C, nepheline and albite buffer log \(a_{{{\text{SiO}}_{{ 2 , {\text{aq}}}} }}\) to equilibrium values of −4 to −3, and at \(a_{{{\text{H}}_{ 2} {\text{O}}}}\) = 0.5, analcime is stable as the only zeolite mineral species. With decreasing temperatures to 150 °C, natrolite becomes stable at the expense of nepheline at the lower \(a_{{{\text{SiO}}_{ 2} }}\) limit of the analcime stability field. At low-temperature conditions (50 °C), natrolite is the dominant zeolite species, whereas the \(a_{{{\text{H}}_{ 2} {\text{O}}}}\) limit for natrolite stability decreases (Fig. 15). The calculated saturation of quartz for the respective conditions plots at higher silica activities than the observed phase equilibrium that buffer the activities of water and SiO2. This is in agreement with the absence of quartz in the late-stage assemblages used for calculating Fig. 15. However, quartz appears locally in association with calcite as breakdown in product of wollastonite. In this case, quartz may be formed as a result of local, small-scale (<1 cm) chemical fluid variations buffered by local breakdown reactants.

The zeolite mineralogy shows a general shift from Ca–Na zeolite species (gonnardite, thomsonite, mesolite) to the Na-endmember natrolite. The chemical evolution of fluids during zeolite formation can be expressed by the observed sequence marked by Ca–Na zeolite species and natrolite.

Figure 14 illustrates the change in fluid composition during the hydrothermal replacement as a function of aqueous cation (Ca2+, Na+) to hydrogen ion activity ratios at temperatures of 50 and 100 °C, and pressure of 10 MPa. The overall topology of stability fields does not change with small variations of pressure; therefore, the uncertainty of the emplacement depth can be ignored. Gonnardite does not appear in the stability diagram due to a lack of reliable thermodynamic data for this phase, instead mesolite appears as intermediate Ca–Na zeolite (Rogers et al. 2006). The zeolite paragenesis gonnardite–natrolite corresponds to a decrease in the Ca/Na ratio, as well as an increase in the Si/Al ratio with time (Figs. 10, 11). The hypothetical fluid evolution path during zeolite formation in the Fohberg phonolite shows an increase in Na+ activity relative to Ca2+ activity. A temperature or pressure change causing the change in mineralogy is implausible according to Fig. 14. Passaglia and Sheppard (2001) stated that unusual crystallization conditions (high temperature and H2O pressure) favor the formation of disordered gonnardite–tetranatrolite instead of ordered natrolite.

Na–Ca feldspathoid minerals and wollastonite breakdown releases elements necessary for zeolite formation. We assume that the primary feldspathoid minerals are of mixed Ca–Na composition. Early precipitation of Ca–Na zeolites species occurred at low log[\(a_{{{\text{Na}}^{ + } }}^{{}}\))/(\(a_{{{\text{H}}^{ + } }}^{{}}\))]. With time, the fluid increased in log[(\(a_{{{\text{Na}}^{ + } }}^{{}}\))/(\(a_{{{\text{H}}^{ + } }}^{{}}\))] and decreased in log[(\(a_{{{\text{Ca}}^{2 + } }}^{{}}\))/(\(a_{{{\text{H}}^{ + } }}^{2}\))] and moved into the stability field of natrolite (Fig. 14). This change in fluid composition is host rock controlled by a time sequence in primary mineral alteration, from early wollastonite breakdown to a later feldspathoid mineral dominated breakdown.

Similar sequences (gonnardite–natrolite) are observed in basaltic lavas in the Disko–Nuussuaq region, West Greenland (Neuhoff et al. 2006; Rogers et al. 2006) and in the Kahrizak area, Iran (Kousehlar et al. 2012), and are interpreted to have formed in a chemically distinct alteration style that reflects the less Ca- and Si-rich primary compositions of these lavas in contrast to provinces with more evolved basaltic rocks, like in Iceland (Neuhoff et al. 1999; Weisenberger and Selbekk 2009) or East Greenland (Neuhoff et al. 1997).

Calcium source

Due to the strong chemical affinity of Sr and Ca, the 87Sr/86Sr ratios can be used to get information about the calcium source in secondary calcite (Land 1987; Schultz et al. 1989). The strontium isotopic composition of fracture calcite and alteration assemblage calcite in the Fohberg phonolite is in the same range than the volcanic rocks of the KVC (Schleicher et al. 1990) (Fig. 13). This suggests that calcium is locally derived from the volcanic rocks. The strontium isotopic composition of alteration calcite is significantly lower than the strontium ratio of seawater during Mesozoic and Cenozoic times (Burke et al. 1982). This indicates that dissolution of marine carbonates in the sedimentary pile of the eastern Kaiserstuhl does not constitute a significant source of strontium and, by proxy, calcium. Similar results are found for the Pleistocene loess cover of the KVC. Strontium isotopic compositions of bulk loss samples (Wimmenauer 2010) are more radiogenic than calcite in the Fohberg phonolite (Fig. 13) and loess can be excluded as major source. Wimmenauer (2010) described calcite-filled fractures in volcanic rocks of the KVC. For calcite-filled fracture without any obvious source other than volcanic rocks of the KVC, he observed similar strontium isotopic ratios than the KVC rock (Fig. 13). Higher radiogenic ratios for calcite-filled fractures hosted in a limburgitic lava flow near Sasbach and tephritic lava flows near Achkarren (Fig. 13) are consistent with petrographic and local geological observations that indicate external sedimentary contributions.

Comparing alteration and fracture calcite, fracture calcite shows slightly higher radiogenic strontium ratios than corresponding alteration calcite (Fig. 13; Table 5). We suggest that alteration calcite received its calcium from primary igneous phases in close vicinity (cm to dm scale). The slightly different 87Sr/86Sr ratios of alteration calcite in different samples indicate local heterogeneities of strontium isotopic compositions within the phonolitic intrusive body. The higher strontium ratio for fracture calcite may be caused by calcium derived from surrounding sedimentary units. The significant differences in 87Sr/86Sr ratios of different fracture calcite samples indicate a heterogeneous fluid in the more open fracture system. Sample Sp12-Cc02 is located in the core of the intrusion and has a lower strontium ratio than Sp12-Cc05, which is derived from the outer shell of the intrusive body. This suggests that the outer parts of the intrusive body had a stronger interaction with fluids derived from the surrounding sedimentary sequence than the core of the intrusion.

Fluid source

The oxygen and carbon stable isotope compositions of calcite can be used to determine the source characteristics of the late-stage fluids. Information on early fluids that caused major alteration and zeolite formation cannot be assessed, due to the lack of zeolites suitable for stable isotope geochemistry. However, a H2O-dominated fluid with low CO2 tolerance at elevated temperature (~100 °C) is most likely, in comparison with the formation of natrolite in similar environments (Kousehlar et al. 2012).

For late calcite in fractures, two different groups can be classified according to their δ13CVPDB and δ18OSMOW values (Fig. 12; Table 5). The first group is homogenous and clusters at δ18OSMOW values between 23 and 25 ‰ and δ13CVPDB values between −13 and −10 ‰ and is similar to secondary calcite in volcanic rocks of the KVC described by Hubberten et al. (1988) (Fig. 12). The light δ13CVPDB isotopic signature could be caused by disintegrated organic matter in sediments of the Paleogene Pechelbronn formation, which are known for their organic content elsewhere in the Upper Rhine Graben (Person and Garven 1992). This is in agreement with the appearance of bitumen as late-stage fracture fill described by Albrecht (1981). A similar late-stage evolution of peralkaline rocks is found in the Mont Saint–Hilaire complex (Schilling et al. 2011). By using zeolite stability, calcite precipitation can be assumed to have occurred below 100 °C. Applying the fractionation equation of Friedman and O’Neil (1977) and assuming temperatures for calcite formation between 25 and 100 °C, fracture calcite δ18O values constrain the δ18O composition of the precipitating pore fluid to −6 to +6 ‰SMOW. Ascending deep geothermal fluids like those found elsewhere in the Upper Rhine Graben (δ18OSMOW: −10 to −1 ‰ at 120–150 °C, Pauwels et al. 1993) would cause calcite with a significantly lighter δ18OSMOW signature and are an unlikely source for calcite found in fractures of the Fohberg phonolite. More likely are meteoric pore fluids whose compositions are controlled by weathering reactions of volcanic rocks (Hubberten et al. 1988). The second group shows a negative correlation between δ13CVPDB and δ18OSMOW (Fig. 13). These calcite samples are several mm-sized euhedral late-stage crystals growing on top of an earlier fracture generation (Sp12-Cc02, Sp12-Cc03) and as layered fibers perpendicular to the fracture wall (Table 5), reflecting late-stage calcite mineralization. The compositional range of these calcites can be explained by low-temperature reactions and isotopic exchange with later bicarbonate-dominated meteoric water. Remobilization of calcite through interaction with such bicarbonate-dominated meteoric water will shift the δ13C value of precipitating secondary calcites to heavier values (Schwinn et al. 2006). Marine Jurassic limestones and Paleogene sediments that are exposed on the eastern part of the KVC may buffer this later bicarbonate-dominated meteoric water.

Fluids derived from the overlying Pleistocene loess unit can be excluded as fluid source. This is in agreement with observation by Wimmenauer (2010) in calcite-filled fractures in tephritic rocks of the KVC. Wimmenauer (2010) stated that the calcite fracture-filling phase formed most probably during the younger Tertiary and well before the deposition of the loess cover in the Pleistocene.

Conclusions

Our investigation of the subsolidus hydrothermal evolution of the Fohberg phonolite intrusion yields the following results and is summarized in Fig. 16. After emplacement of the phonolite magma, primary minerals crystallized from the melt. Crystallization and cooling continued and phase separation and boiling may have led to the intense fracturing of the phonolite body.
Fig. 16

Schematic figure illustrating the evolution of the Fohberg phonolite. a Sequence of processes affecting the cooling intrusive body emplaced in sedimentary units. b Magmatic and subsolidus (alteration) assemblages of major minerals as indicated by textural observations

Syn- to post-fracturing fluid-driven re-equilibration of feldspathoid minerals and wollastonite causes the breakdown of the primary phases and formation of a set of secondary phases. Feldspathoid minerals are totally replaced by secondary phases including various zeolite species, calcite, and barite. Precursor feldspathoid minerals are most likely haüyne and nepheline due to the shape of phenocryst pseudomorphs and the appearance of sulfate minerals as secondary phases, in particular barite. The breakdown of wollastonite results in the formation of various zeolites, calcite, pectolite, sepiolite, and quartz. The high variability of wollastonite breakdown product suggests a heterogenic fluid composition throughout the Fohberg phonolite intrusion.

Zeolites formed during subsolidus hydrothermal alteration (<150 °C) under alkaline conditions and completely replace feldspathoid minerals in the matrix of the rock. Assuming a temperature of zeolite formation below 150 °C, a fluid with low CO2 tolerance (\(X_{{{\text{CO}}_{ 2} }}\) = <0.01) causes the breakdown of wollastonite and formation of zeolites. The fluid composition is characterized by low log \(a_{{{\text{SiO}}_{{ 2 , {\text{aq}}}} }}\) values. Higher log \(a_{{{\text{SiO}}_{{ 2 , {\text{aq}}}} }}\) would favor the formation of analcime. Zeolite formation is accompanied by a change in mineralogy from Ca–Na-dominated zeolite species, like gonnardite and mesolite to the pure sodium zeolite natrolite. The sequence reflects an increase in log[(\(a_{{{\text{Na}}^{ + } }}^{{}}\))/(\(a_{{{\text{H}}^{ + } }}^{{}}\))] and decrease in log[(\(a_{{{\text{Ca}}^{2 + } }}^{{}}\))/(\(a_{{{\text{H}}^{ + } }}^{2}\))] of the precipitating fluid rather than a change in temperature.

The Fohberg intrusion is crosscut by numerous fractures, which are totally or partially sealed. Fracture minerals show a similar zeolite sequence as observed in the host rock matrix, followed by calcite and a variety of other silicates, carbonates, and sulfates as younger generations. Stable isotope analyses of calcite as late-stage fracture precipitate indicate meteoric fluid circulation. The external meteoric fluids were probably in equilibrium with the surrounding sedimentary rocks, which contribute to the oxygen and carbon isotopic compositions of calcite. This is indicated by the appearance of hydrocarbons, derived from mobilized organic matter of the surrounding sedimentary units. Calcite is remobilized during very late stages.

The 87Sr/86Sr analysis of matrix and fracture calcite reveals low radiogenic values, indicative for primary igneous rocks of the KVC. This indicates the Ca2+ and most probably other elements except H2O are locally derived from breakdown of primary igneous phases.

Notes

Acknowledgments

We are grateful to Hans G. Hauri Mineralstoffwerke for providing access to the Fohberg quarry and support of this study. We thank Sari Forss for thin section preparation, Leena Palmu for her advice and help with the electron microprobe, and Shenhong Yang for help with Sr-isotopic analyses. We thank Lieven Machiels for his detailed and constructive comments and Wolf-Christian Dullo for his efforts and the editorial handling of the paper. Special thanks to the Holopainen Foundation for the financial support.

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Copyright information

© Springer-Verlag Berlin Heidelberg 2014

Authors and Affiliations

  • Tobias Björn Weisenberger
    • 1
  • Simon Spürgin
    • 2
  • Yann Lahaye
    • 3
  1. 1.Department of GeosciencesUniversity of OuluOuluFinland
  2. 2.Hans G. Hauri MineralstoffwerkeBötzingenGermany
  3. 3.Geological Survey of FinlandEspooFinland

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