around the rim of a void.
ii) late; precipitates from evolved or from second generation fluids, that have crystallized in the core of voids.
Examples are illustrated in Fig 2.2.2. 2.2.4.4 Vein systems
At several horizons amygdaloidal assemblages are cut by vertical vein systems bearing less hydrous zeolite species. The majority of the veins penetrate along cooling joints in the basalt.
Several veins display antitaxial features (Durney and Ramsay, 1973), characteristic of syntectonic growth. Identical zeolite species are usually contained in both vein systems. The veins extend upwards from a region where the vein material occurs as the dominant amygdale filling. For example, in the Kilpatricks (Fig 2.1.1):-
i) basalt with amygdaloidal natrolite, is overlain by basalt with amygdaloidal analcime plus vein natrolite. ii) basalt with amygdaloidal analcime, is overlain by basalt
with amygdaloidal quartz and chlorite, plus vein analcime
Likely origins of the vein and amygdale systems are discussed below,
2.2.4.5 Porosity and Permeability
Amygdale and vein assemblages reflect initial rock porosity and permeability. For example, it is suggested in
a) chapters 2.2 and 2.3; that the extent of alteration was related to variations in permeability/porosity within individual flows or flow units
zation at shallower burial depths than in basalts.
Vein systems
In the metavolcanics of the Midland Valley less hydrous (higher temperature) zeolites have been found to be stable in veins, whilst more hydrous^ phases (lower temperature) occur in amygdales. For example, at Burntisland (chapter 2.2), natrolite occupies veins whilst analcime occurs in amygdales at the same stratigraphic level.
The origin of the veins is problematical, since vein assemblages occur at all levels within the volcanics. Consequently, variations in Pioad or variations in bulk composition between the vein and the adjacent amygdaloidal region, are discounted as a possible cause of the anomally.
The distribution of vein and amygdale assemblages may be accounted for in three ways:-
i) As a result of temperature variations. This would require that vein minerals were precipitated from metamorphic
'••• fluids which were, either driven-off from intrusive bodies or were liberated during dehydration reactions at greater depths. The former, was envisaged by Walker (1960b) to account for zeolitization in the basalts of eastern Iceland. Under such conditions, fluids in amygdaloidal regions would have been intergranular; hence Pjith ~ ^fluid*
ii) As a result of higher fluid/rock ratios in the veins. According to Houghton (1982), this could lead to chemical differences in the fluids of the two environments.
iii) As a result of variations in Pu A during metamorphism, due
h2u
to chemical dilution of the fluid phase. Thus whereas Pfiuid = pload during metamorphism, P„ n may be less than Pfiuid* The
h2u
the most likely varient would have been the differing chemical potential (u) of £/C02//jH20
The chemical potential {or molar free energy) of an ideal gas is related to pressure (P) by the
following equation:-
M = W0 + RT in P.
Where ju® = ju at unit P R = gas constant
T = absolute temperature
However it is unlikely that variable £iC02/juH20 ratios were responsible for variations in Pfiuid in the Midland Valley. It was demonstrated in chapter 2.2 that there were a sparcity of clay-carbonate assemblages in the zeolite sequences. Thompson (1971a) suggests that this reflects a consistently low value of ju C02 throughout amygdular hydrothermal alteration.
Point (i) has established that fluid (?) in amygdular regions was under lithostatic pressure (Puth = pfluid^* However, Houghton (1982) demonstrated that at shallow burial depths, major fractures are open to the surface, so that fluid contained in veins was under hydrostatic pressure (Phyd = x/3 plith)* Thus higher fluid pressures
in amygdales may have favoured the formation of more hydrous species, under conditions where less hydrous phases were stable in veins.
The factors responsible for crystallization can be derived by comparing equations for the chemical potential of vein and amygdale mineral growth under hydrostatic and lithostatic pressures,
respectively. A general equation for a vein crystal growing in a fluid under hydrostatic pressure (P^), as determined by Everett
(1961) is given below:-
U = u (Ph) + Vs o (SA/SV)
where Vs = molar volume of solid
o = interfacial energy between solid and liquid A = area of interface
V = volume of solid
This may be compared with Yardley’s (1975) equation for crystal growth under lithostatic (PP) conditions.
/j=ju + w + Vso (SA/SV) + E
where w is dependant upon surface tension (o), local stress deviation, and A P^ - Pp, E = energy,
In the latter instance, pore fluids are intergranular, and hence the surface of all the crystals will be in equilibrium with the adjacent fluid. Movement of material into the vesicles will therefore have resulted from a potential
(v)
gradient operating between the vesicle and the groundmass fluid. This was defined byYardley (1975), as:*
$/<{/“ ( A^g, * A^y) /X
where a
^
= y of groundless grain Mv ’ M of vein grainx = distance between grains
However, considering that zeolitization requires progressive dehydration reaction, liberated water must have moved along grain boundaries and generally upwards to lower Pfiui<p Any vesicles would have been filled by fluid under confining conditions of Pfiui<i = ^lith •
1 Fluid motion would have been governed by:-
i) diffusion through crystal lattices ii) interstitial or porosity flow iii) fracture flow
Petrographic evidence for diffusion is sited in chapter 2,3. Briefly: flow was concentrated in fractures; alteration haloes were developed in the adjacent metabasalts and reflected exchanging fluid: rock reactions, (see chapter 2.6 for further discussion). The chemical potential
(
m)
of the two areas would have been different, therefore material would have diffused from the vein to the wall-rock to eliminate the potential (ju) gradient. Fyfeet
si, (1978), define the flux (J) of material (i) across a plane perpendicular to thedirection of mass transfer as:
= D| (dc|/dx)t cm2 s’1
Where dc/dx = a concentration gradient. This is proportional to a potential gradient (dju/dx) across the plane.
where D| = the diffusion co-efficient t = time
x = mole fraction of component i
The equation shows that flow was controlled by the chemical potential of the fluid and by the length of time that channel-ways were open. By inference, zeolitization (although self-propagating exothermic reactions) must have been controlled by the relative rates of fluid access.
Interstitial/porosity flow occurred through macro-pore (vesicles) and micro-pore (along intergranular boundaries) spaces. The effectiveness of the process varies as a function of:-
i) pressure; this reduces grain boundary volume through compaction
ii) time; progressive sealing of pore spaces due to precipitation from fluids.
Both variables reduce pore volume (porosity) and the degree of
Table 2.2.3
Rock Type Region k(cm2) Reference
Dense rhyolite Vesicular rhyodacite Opsidian Dane, California IO-8 to IO-12 0.2 x io”8 Eichelberger et al.t 1986 Vesicular basalt Vesicular basalt Oahu, Hawaii
Wairakei, New Zealand H-*
H 1 o o 1 1
f
O ' O Cheng & Minkowycz, 1977 Bedded tuff Welded tuff unspecified unspecified 10_1° to IO’12 <2 x IO-16Norton & Knapp, 1977
interconnection between pores (permeability).
Fracture flow reflects fluid movement along veins under the physiochemical conditions to be established in section 2. For hydraulic fracture to occur however, must exceed the tensile strength (o3) of the rock.
The total porosity of the rock (Norton and Knapp, 1977) is a function of fracture flow, diffusion porosity and residual pore space. Permeability, however is governed by the abundance and geometry of continuous flow channels. Measurements of permeability (k) must also take into account the temperature (T) and the viscosity (n) of the fluid involved, such that:-
k = (Q/A)n(dp/ds)~1 after Fyfe et al, 1978
Where Q equals the volume of fluid which passes with time through an area (A) in the direction of the pressure gradient (dp/ds). Measure ments are quoted in Darcies D, where ID = 1 cm3 s"1 cm”2 at 1 bar cm"1; T = 20 °C.
Representative permeability measurements in certain volcanic rocks are presented in table 2.2.3. These values may be compared with permeability measurements in fractures. Toulmin and Clark (1967)
considered that in a dense rock, (where k
s
10”5D), k = 3 x 103D in a 1 ju fracture and Ks
3000 D in a 1 mm fracture. The data suggests that fractures significantly increase the flow of hydrothermal fluids in volcanics.2.2.4.6 Amygdale assemblages of the eastern Midland Valley Amygdale assemblages are developed in four geographically
distinct regions (Fig 2.2.1) of the eastern Midland Valley
i) Bathgate
ii) Burntisland (and East Fife)
iii) Arthur’s Seat and Craiglockhart Hill, Edinburgh iv) the Garleton Hills
Sampled localities are shown on figure 2.2.1, and are listed in appendix 2.2.1. Zeolite distributions are presented in figure 2.2.3. Eight of Heddle’s (1924) localities are incorporated on the map; the
remainder represent new zeolite occurrences in the Midland Valley. The following paragraphs describe the amygdale assemblages from the regions listed above. Explanations of the small scale features are given. The overall significance of the amygdale assemblages to hydrothermal alteration will be discussed later in the chapter.
Bathgate
Evidence for hydrothermal alteration in the Bathgate Volcanics is limited. Basalts are generally fresh, although green ’spots’ related to chlorite crystallization (confirmed by microprobe analyses - chapter 2.4) may occur.
Vesicular and/or amygdaloidal horizons are present; neither are abundant. Amygdales and veins contain quartz and/or calcite. Vein mineralization (at CV22 and CV24) can be directly related to the proximity of quartz-dolorite sills and dykes. Since amygdale horizons are rare, it is impossible to determine whether their contents