Excess argon in K–Ar and Ar–Ar geochronology

Review article

Excess argon in K–Ar and Ar–Ar geochronology

Simon Kelley *

Department of Earth Sciences, Open University, Walton Hall, Milton Keynes, MK7 6AA, UK

Received 13 December 2000; accepted 9 April 2002

Abstract

The K–Ar and Ar–Ar dating techniques occasionally produce anomalously old ages attributed to excess argon, and such

data is often rejected as not offering any insight into the age, thermal history or geochemistry of the rock. However,

improvements in the quantification of argon geochemistry now provide a framework to model excess argon in both open and

closed systems. Solubility data for argon in hydrous fluids, melts and emerging data for minerals can be used to understand the

behaviour of excess argon, and provide valuable insights into the environment in which the samples cooled to their argon

retention or ‘closure’ temperature. Treating excess argon as a trace element also throws light on its behaviour in minerals above

the closure temperature, in deeply buried dry systems such as eclogites, blueschists, granulites and even in the lithospheric

mantle. Extremely low partition coefficients between K-feldspar and hydrous fluid phases predict lower excess argon

susceptibility than micas and this is observed in fluid-poor systems. Variation of partition coefficients can lead to excess argon

in fluids being introduced into minerals or removed from minerals as grain boundary fluids change during flow through a rock.

However, excess argon can also be introduced or removed from minerals by varying temperature, without the need for fluid

flow. High mineral/melt and mineral/fluid partition coefficients are also the reason why excess argon is often concentrated in

inclusions within minerals. Partition coefficients between minerals and hydrous fluids as low as 10� 6 lead fluid inclusions to

dominate the radiogenic argon budget, particularly in low potassium minerals. Melt inclusions are less dominant but become

critical in dating younger samples.

D 2002 Published by Elsevier Science B.V.

Keywords: Excess argon; K–Ar and Ar–Ar geochronology; Trace element

1. Introduction

Most isotope geochronometers, such as Rb–Sr,

include a measurement of initial radiogenic daughter

product concentrations, but the simple assumption

often made in order to calculate K–Ar and some

Ar–Ar dates is that the samples initially contained

no radiogenic argon. Argon escaping from minerals

above their closure temperature is assumed to enter an

‘infinite reservoir’, leaving minerals free of radiogenic

argon. The dilemma in this scenario is that studies of

ground waters, fluid inclusions and many natural

volcanic glasses demonstrate that hydrous fluids and

magmas at all depths in the crust contain argon.

Moreover, most if not all of these fluids and magmas

contain argon enriched in the radiogenic isotope 40Ar,

in other words they contain excess argon. Following

this line of logic, it seems evident that all minerals in

the crust are bathed in fluids containing excess argon,

so why does K–Ar dating work at all? Why don’t all

0009-2541/02/$ - see front matter D 2002 Published by Elsevier Science B.V.

PII: S0009 -2541 (02 )00064 -5

* Tel.: +44-1908-653009.

E-mail address: [email protected] (S. Kelley).

www.elsevier.com/locate/chemgeo

Chemical Geology 188 (2002) 1–22

samples contain at least small amounts of excess

argon?

First we must define excess argon. When the con-

centration of 40Ar in minerals exceeds that generated

by in situ decay of potassium, the phenomenon is

called ‘extraneous argon’ and can have several causes.

Contamination by older grains such as soils, which

become incorporated into ignimbrite eruptions, is

commonly known as ‘inherited argon’. There are

other less clear cut examples of inherited argon such

as partially reset minerals in metamorphic rocks, as

we will see later, but in all cases inherited argon has a

source within the system. The other main reason why

radiogenic 40Ar contents exceed in situ production is

the introduction of ‘excess argon’ from outside the

system. This review will consider ‘excess argon’ and

cases where a closed system blurs the boundary

between ‘excess argon’ and ‘inherited argon’ in

fluid-poor systems (McDougall and Harrison, 1999;

Scaillet, 1998).

Argon is a trace element and is present in all rocks

at trace levels, though for many geochronological

purposes, it can be ignored. However, it is always

present and can reach concentrations that cause a K–

Ar or Ar–Ar age to be significantly older than the

event which initialised or reset the isotope system. K–

Ar and Ar–Ar ages are often corrected for argon with

an atmospheric ratio (295.5, by convention, Steiger

and Jager, 1977), but this ratio is only observed in

nature in the atmosphere and in surface waters.

Natural basin brines, metamorphic fluids and melts

do not contain argon with an atmospheric ratio, they

all exhibit an excess of 40Ar.

The first reports of excess argon appeared soon after

K–Ar dating had been established in the 1950s. Damon

and Kulp (1958) found excess argon in beryl, cordierite

and tourmaline but this was soon followed by reports of

excess argon in olivine and whole rock basalt (Dal-

rymple and Moore, 1968), feldspar (Livingston et al.,

1967), mica (Lovering and Richards, 1964), and

amphibole (Pearson et al., 1966). However, possibly

the most important early study of excess argon was that

of Rama et al. (1965) who measured large quantities of

excess argon in fluid inclusions within quartz and

fluorite. Early K–Ar literature sometimes used the

term excess argon to mean any age older than expected,

and there was also confusion over the incorporation

mechanisms for excess argon, and a debate over

whether excess argon was occluded during crystallisa-

tion or subsequently diffused into minerals, (cf. Dal-

rymple and Lanphere, 1969). This confusion stemmed

from lack of data on either argon diffusion rates or the

concentrations of excess argon in fluid inclusions

within minerals. More recently, excess argon has come

to mean parent-less radiogenic argon incorporated into

a mineral during crystallisation, introduced into the

mineral lattice by subsequent diffusion or occluded

within fluid or melt inclusions within the mineral. This

definition specifically excludes inherited argon caused

by incorporation of older grains in deposits such as

tuffs or partially reset metamorphic rocks.

Excess argon also appeared in early Ar–Ar meas-

urements (e.g., Pankhurst et al., 1973), though the use

of stepped heating also afforded the opportunity to

discriminate against excess argon by removing the

need to assume an atmospheric initial ratio for con-

taminating argon within samples. By making use of

the isochron technique (Roddick, 1978), many suc-

cessful age determinations have been made in systems

with excess argon (e.g., Heizler and Harrison, 1988).

In fact many published Ar–Ar ages contain small

amounts of excess argon reflected in an initial40Ar/36Ar ratio which is within a few percent of the

atmospheric ratio (40Ar/39Ar = 295.5, Nier, 1950).

Such determinations yield precise results when the40Ar/36Ar ratio of the contaminant is close to that of

atmospheric argon. However, as the initial ratio

increases and the 36Ar peak becomes more difficult

to measure, the precision of resultant ages is quickly

compromised (e.g., Arnaud and Kelley, 1995; Sher-

lock and Arnaud, 1999). In addition, the initial ratio

correction only produces precise ages when the iso-

tope ratio of the contaminating component is homo-

geneous. If the sample contains fluid inclusions with

variable 40Ar/36Ar ratios for example, the scatter of

data prevents precise age determination (e.g., Cumb-

est et al., 1994; Reddy et al., 1997).

The development of new analytical techniques for

Ar–Ar dating has resulted in many insights into the

behaviour of excess argon. Ar–Ar stepped heating

provided a physical technique to separate and analyse

phases within individual samples as a result of their

different breakdown temperatures (e.g., Hanes et al.,

1985; Belluso et al., 2000). Stepped heating also

produces decrepitation of fluid inclusions at low

temperatures resulting in the high initial ages com-

S. Kelley / Chemical Geology 188 (2002) 1–222

monly observed in release spectra. Stepped heating

has also demonstrated excess argon diffusion into

grain boundaries (Harrison and McDougall, 1981),

and multiple excess argon components in mineral

separates (Heizler and Harrison, 1988). One feature

of stepped heating samples containing excess argon is

the saddle or ‘U’-shaped Ar–Ar stepped heating

release spectrum (Lanphere and Dalrymple, 1976;

McDougall and Harrison, 1999; Wartho et al.,

1996), commonly associated with low potassium

rocks and minerals such as plagioclase, amphibole

and clinopyroxene. Several explanations have been

offered for this release pattern, first described by

Dalrymple et al. (1975) and Lanphere and Dalrymple

(1976) in samples from dykes intruding Pre-Cambrian

rocks in Liberia (Fig. 1A). The initial Ar release yields

high apparent ages, which decrease with progressive39Ar release, approaching the true age of the sample,

and finally returning to high values towards the end of

argon release. Although arguments have been made

for excess argon incorporation via special diffusion

mechanisms such as anion diffusion, (Harrison and

McDougall, 1981), recent data seem to demonstrate

that the most likely candidates are inclusions. Melt or

fluid inclusions release argon at low temperature, and

melt, fluid or solid inclusions, release argon at high

temperature as minerals break down or melt at high

temperature (Esser et al., 1997; Boven et al., 2001).

The reasons for high concentrations of excess argon in

inclusions are considered below. Similar patterns have

been seen in K-feldspars and these also probably

relate to the interplay of fluid inclusion decrepitation

at low temperature and excess argon in large sub-

grains (or solid inclusions) released at high temper-

ature (Zeitler and FitzGerald, 1986; Harrison et al.,

1994; Foster et al., 1990).

In vacuo crushing has provided another technique

to study the close correspondence between excess

argon and saline crustal fluids trapped in fluid inclu-

sions in quartz (Kelley et al., 1986; Turner and

Bannon, 1992; Turner et al., 1993) and K-feldspar

(Turner and Wang, 1992; Burgess et al., 1992; Harri-

son et al., 1993, 1994). In vacuo crushing very

effectively separates argon in fluid from argon in the

solid, and Cl concentrations derived from the 38Ar

peak (cf. McDougall and Harrison, 1999) can be

combined with salinity measurements to monitor

argon trace concentrations and isotope ratios in the

fluid. Finally, in situ laser spot extraction techniques

have provided a method of investigating excess argon

distributions within minerals such as phlogopite (e.g.,

Phillips and Onstott, 1988; Phillips, 1991) and phen-

gite (e.g., Boundy et al., 1996, 1997; Scaillet et al.,

1992; Scaillet, 1996). Laser studies have also been

able to demonstrate a close correlation of excess argon

with fluid mediated alteration in ultra-high pressure

UHP terrains (e.g., Giorgis et al., 2000), diffusion of

excess argon through the mineral lattice (e.g., Lee et

al., 1990; Reddy et al., 1996; Pickles et al., 1997)

(Fig. 1B) and incorporation in solid inclusions (Boven

et al., 2001).

Fig. 1. Natural examples of excess argon. (A) A stepped heating

‘saddle’-shaped 39Ar release spectrum produced by a plagioclase

separate from a dyke in Liberia (Lanphere and Dalrymple, 1976).

(B) Excess argon diffusion measured parallel to biotite cleavage in a

rock from the IIDK, western Alps (Pickles et al., 1997).

S. Kelley / Chemical Geology 188 (2002) 1–22 3

A brief survey of the list of examples above shows

that studies of excess argon have tended to be con-

centrated in metamorphic studies and this reflects the

distribution of excess argon in natural samples. Excess

argon is less common in volcanic systems where

outgassing provides a release mechanism. However

in hyperbyssal and plutonic rocks excess argon is

more common. Excess argon is also particularly

common in hydrothermal systems associated with

large granite intrusions (Kelley et al., 1986, Turner

et al., 1993), shear zones in ancient metamorphic

terrains (Allen and Stubbs, 1982, Smith et al., 1994,

Vance et al., 1998), contact metamorphic aureoles

(McDougall and Harrison, 1999), and ultra-high pres-

sure metamorphic rocks (Arnaud and Kelley, 1995;

Inger et al., 1996; Scaillet, 1996; Reddy et al., 1996;

Li et al., 1994; Ruffet et al., 1997; Sherlock et al.,

1999; Sherlock and Arnaud, 1999).

2. Argon geochemistry

Excess argon is generally discussed on a case-by-

case basis, but by considering the whole system

including the grain boundaries, as we would for any

other trace element, it is possible to construct a

framework which provides a more powerful way to

describe excess argon. The fact that argon is only

present in trace amounts and is inert, means that as a

first-order assumption, we can discuss its distribution

between the phases in any system in terms of sol-

ubility. In most cases this is expressed as a solid/melt

or solid/liquid partition coefficient, and the develop-

ment of excess argon in any system can be modelled

in terms of exchange between phases (including grain

boundary fluid) and build-up of radiogenic daughter40Ar. It is difficult to measure argon solubility and

partition coefficients in nature, but laboratory studies

allow us to model the distribution of argon in natural

systems. Argon partition coefficients are shown as KD

in the present work, rather than DAr, to prevent

confusion with the argon diffusion coefficient. In

order to construct a model for excess argon develop-

ment, we first we need to review the data on argon

solubility in hydrous fluids, melts and minerals. In

addition, weight concentrations (ppm and ppb) are

adopted in this work for comparison with other trace

elements.

2.1. Argon solubility in water

Fundamental to all considerations of excess

argon is the amount, salinity and source of hydrous

fluid in the system, whether it is in cracks, pore

filling, grain boundary fluid or even single molec-

ular layer coating grains. It is this fluid which in

most cases represents the ‘infinite reservoir’ into

which 40Ar* passes from mineral grains and is lost

from the rock system. Laboratory studies have

provided a basis for calculating argon solubility in

natural water systems (Weiss, 1970; Crovetto et al.,

1982; Smith and Kennedy, 1983). The concentra-

tion of argon in all surface and near surface waters

is dominated by equilibrium between surface water

and the atmosphere. Moreover, studies of sub-sur-

face ground waters and basinal brines show argon

concentrations at least as high as surface waters. In

fact, argon acts as a conservative tracer in near

surface environments where it has been used to

calculate crustal degassing rates (Torgersen et al.,

1989) and the temperatures of waters entering

aquifer systems at source (Mazor, 1972; Ballentine

and Hall, 1999). At deeper levels, fluid inclusions

in hydrothermal systems (Kelley et al., 1986; Ken-

drick et al., 2001a,b; Turner and Bannon, 1992;

Turner et al., 1993), contain argon at similar or greater

concentrations than surface waters. Thus, in near sur-

face environments, hydrous fluids appear to contain

atmospheric argon at roughly similar levels to river

and seawater with additional argon derived from the

host rocks.

The solubility of argon in pure water has been

characterised by several workers (Weiss, 1970; Cro-

vetto et al., 1982), who showed that argon solubility

follows Henry’s law with a constant (kH) which varies

with temperature (Fig. 2). Argon solubility is often

displayed as a Bunsen coefficient (the concentration

of argon in cm3 STP/l water at equilibrium with argon

at 1 atm). However, in the present work the equili-

brium concentration in pure surface water is illus-

trated as ppm in water in equilibrium with 1 bar

pressure of argon, for comparison with later descrip-

tions of argon solubility in melts and minerals (Fig.

3). The amounts of argon found in pure water in

equilibrium with the atmosphere can easily be calcu-

lated from Fig. 3 by multiplying the solubility by

0.00934, the proportion of argon in air.


Fig. 3. Variation of argon solubility in water in equilibrium with argon at 1 bar pressure (in ppm) with temperature and salinity. The solid line is

the data for pure water (Weiss, 1970). Short dashes are data from Crovetto et al. (1982) for water at high temperature, long dashes are data for

water with 13% and 26% NaCl, from Smith and Kennedy (1983).

Fig. 2. Variation of the natural log of Henry’s constant with temperature and salinity. The solid line is the data for pure water (Weiss, 1970).

Short dashes are data from Crovetto et al. (1982) for water at high temperature, long dashes are data for water with 13% and 26% NaCl, from

Smith and Kennedy (1983).


The great majority of groundwater is saline and

exhibits decreasing solubility with increasing NaCl

content. Smith and Kennedy (1983) measured the

‘‘salting out’’ of argon, i.e. the relationship between

the NaCl load and decrease in argon solubility. They

found that in highly saline brines, argon solubility was

reduced by a factor of 3 (Figs. 2 and 3). Some data are

also available on the salting out effect of other solutes

(Smith and Kennedy, 1983 and references therein).

The measured noble gas solubilities have successfully

been used to study noble gas concentrations in saline

groundwaters and predict recharge temperatures (e.g.,

Mazor, 1991; Ballentine and Hall, 1999).

The narrow range of temperatures analysed by

Smith and Kennedy (1983) represent waters close to

those at the Earth’s surface, not pore waters in rocks

above mineral closure temperatures. However, Cro-

vetto et al. (1982) extended the known solubility

range of argon in pure water to 295 jC, showing that

trends at lower temperatures extended to higher tem-

peratures and that argon solubility rises from a mini-

mum (a maximum in kH) at circa. 90 jC (Figs. 2 and

3). Although the extended dataset shows that

increased temperatures in groundwaters might lead

them to become oversaturated in argon at elevated

temperatures (after equilibrating with atmospheric

argon at the surface), many groundwaters actually

contain small excesses of atmospheric argon (over

their surface values) as a result of equilibration with

trapped air in sediments (Ballentine and Hall, 1999;

Aeschbach-Hertig et al., 2000). Exceptions to the rule

of surface argon concentrations are found in hydro-

thermal systems when boiling has occurred (Mazor et

al., 1988), where argon partitions strongly into the

vapour phase reducing argon concentrations in the

fluid phase.

Atmospheric argon concentrations are also greatly

reduced in hydrous fluids in deeper metamorphic

environments (e.g., Cumbest et al., 1994). The reduc-

tion of atmospheric argon concentrations probably

occurs as a result of fluid expulsion during burial

and the accompanying formation of authigenic clay

minerals. During this process, argon will partition

between the coexisting fluid and clay minerals and

although there is currently no experimental data, the

measured atmospheric contents of clay minerals

(Aronson and Hower, 1976; Clauer and Sambhu,

1995) and atmospheric argon contents of ground

waters indicate partition coefficients in the region of

10 � 2. Subsequent transient metamorphic fluids

formed by dehydration reactions of the clay minerals

will contain a small fraction of the atmospheric argon

concentrations in near surface waters. The loss of

surface argon exacerbates the excess argon problem

since the 40Ar/36Ar ratio of the fluid quickly becomes

more enriched in 40Ar* contributed from crustal rocks

enriched in 40K. Thus, the near atmospheric isotope

ratios of ground waters are replaced by progressively

more radiogenic isotope ratios in fluids at depth,

reflecting the integrated age and potassium content

of the rocks.

2.2. Argon solubility in melts

In contrast to argon solubility in water, argon

solubility in silicate melts is less temperature depend-

ent except for silica glass at low temperature (Carroll

and Stolper, 1991), but reflects strong compositional

dependence.

Noble gas solubility and diffusivity have been

measured in a wide range of melt compositions using

microprobe techniques and glass samples subjected to

high noble gas pressures at elevated temperatures

(Carroll and Stolper, 1993 and references therein).

Solubility varies in the range 0.05–0.8 ppm bar� 1,

but may reach up to 1.8 ppm bar� 1 in pure silica

melt. Carroll and Stolper (1993) showed that ionic

porosity (the ratio of the unit cell volume and the

calculated volume of the constituent anions and cat-

ion) was a better predictor of argon solubility in melts

than either molar volume or density, which had

previously been suggested. However, not all melts

follow the correlation, CaO–MgO–Al2O3–SiO2

(CMAS) melts showed anomalously low solubility,

lower than any other measured melts though their

calculated ionic porosities were within the normal

silicate melt range. Although the relationship between

argon solubility and ionic porosity is only empirical,

argon solubility in common melt compositions can be

characterised by their ionic porosity, which is corre-

lated with silica content. Komatiite has an ionic

porosity of 45.5, basaltic andesite, 47.5, and rhyolite

49.5. Pure silica has an ionic porosity of around 50.5

and a correspondingly high argon solubility. Thus,

melts common to geological environments in the

compositional range basalts to rhyolites have argon


solubility ranges from around 0.05 to 0.1 ppm bar� 1

for komatiite to basalt, and up to 0.8 ppm bar � 1 for

rhyolite. In summary, argon solubility in melts is 40–

600 times lower than solubility of argon in pure H2O

and 8–140 times lower than the most saline ground

waters. However, note that silicate melts normal exist

only at much higher temperatures than hydrous fluids

and are thus inherently less compatible with argon.

One of the outcomes of this work on noble gas

solubility in melts was that solubility of argon and

other noble gases could be compared with CO2 and

H2O, and some conclusions drawn concerning degass-

ing during magma ascent. H2O occurs as both molec-

ular water and OH groups in melts whereas CO2

occurs only in a molecular form. H2O is over an order

of magnitude more soluble in melts than CO2 which

has a solubility similar to Ar (Carroll et al., 1994). The

corollary of this observation is that CO2 and Ar will

be more readily exsolved relative to H2O, in early-

formed vapour phases during magma ascent and

degassing. This is an important consideration in whole

rock or indeed mineral dating of volcanic rocks by

Ar–Ar and K–Ar and probably explains the low

incidence of excess argon in lavas arising from well-

degassed volcanic systems.

2.3. Argon solubility in minerals

Until quite recently, there was very little quantifi-

cation of argon solubility in minerals. Early attempts

to measure solubility using bulk mineral samples were

dogged by experimental artefacts, particularly absorp-

tion onto mineral surfaces and cracks (Hiyagon and

Ozima, 1982, 1986; Broadhurst et al., 1990, 1992).

These bulk experiments yielded partition coefficients

(KD) ranging over many orders of magnitude and in

several cases indicating compatible behaviour (KD > 1).

Recent work using an ultra-violet (UV) laser micro-

extraction technique has shown that the KD values

obtained from bulk experiments were often several

orders of magnitude higher than the true values

(Brooker et al., 1998; Chamorro et al., 2002).

Harrison and McDougall (1981) made one of the

first attempts at quantitative assessment of excess

argon in a mineral/fluid system, introducing the con-

cept of argon acting as a trace element, partitioning

between phases. They measured argon solubility in

plagioclase using bulk mineral samples exposed to

argon at 1000–1200 jC, and 1 atm (1 bar) determin-

ing solubility to be in the range 4� 10� 5–2� 10� 4

cm3 STP g� 1 atm � 1 (71 to 357 ppb Ar atm � 1). The

problem with this range of values is that it is similar to

argon solubilities in melts (Carroll and Stolper, 1993),

implying that argon is a relatively compatible trace

element. Similar conclusions arose from the crystal/

melt partition work of Hiyagon and Ozima (1982,

1986) and Broadhurst et al. (1990, 1992). However,

all such studies suffer from problems of distinguishing

absorbed, adsorbed and inclusion hosted argon from

argon actually incorporated in the mineral lattice.

Onstott et al. (1991) derived the first argon sol-

ubility value using a laser probe technique by meas-

uring the uptake of 36Ar into biotite during a

hydrothermal experiment to measure argon diffusion.

They did this by encapsulating biotite with water,

which contained argon at normal surface atmospheric

concentrations. Onstott et al. estimated argon partition

coefficients between water and biotite of between 0.03

and 0.003 and derived a solubility in the range 3.6–36

ppm bar � 1 in biotite. These values are also higher

than more recent measurements of argon solubility in

melts and reach the lower levels of solubility for argon

in saline fluids. The problem with this value seems to

lie in the calculation of the total atmospheric argon

available in the capsule during the experiment. Onstott

et al. (1991) calculated the total argon budget of the

experiment assuming it was only available from water

added to the capsule. It seems likely that air (with

0.934% Ar) trapped in the capsule and absorbed on

the walls led to a much larger argon reservoir being

available to the mica during the experiment and the

solubility and partition coefficients should be regarded

as maximum values (Onstott, personal communica-

tion, 2000).

In order to avoid problems of mineral separation

and surface absorption (Carroll and Draper, 1994),

recent studies have been undertaken using in situ

methods such as the electron microprobe used in

earlier melt solubility and diffusion experiments (Car-

roll and Draper, 1994; Roselieb et al., 1997). How-

ever, even when the experiments are run at high Ar

pressures (Roselieb et al., 1997) argon concentrations

in crystals co-existing with melts are commonly

below the detection limit of electron microprobes

(around 25–30 ppm). Despite the detection problems,

Roselieb et al. (1997) were able to place a maximum


constraint upon the crystal– liquid partition coefficient

for quartz in a SiO2 melt at 0.006 corresponding to a

maximum estimate for the solubility of argon in

quartz of 3.75 ppb bar � 1. Brooker et al. (1998)

showed that an in situ technique using a focussed

UV laser could be used to extract argon in situ from

melt and laboratory produced crystals. The extracted

gas was measured using a noble gas mass spectrom-

eter down to concentrations of 0.09 ppb, 28000 times

lower than the detection limit of the electron microp-

robe. Brooker et al. (1998) measured crystal–liquid

partition coefficients as low as 0.013 for olivine in

basaltic melt and 0.0016 for clinopyroxene in basaltic

melt. These reflect absolute solubility of 0.09 ppb

bar � 1 for clinopyroxene and 1 ppb bar � 1 for olivine.

Recent work has expanded this to show that clinopyr-

oxene/melt partition coefficients remain stable over

very large pressure and compositional ranges (Cha-

morro et al., 2002). Further estimates are available for

fluoro-phlogopite (Roselieb et al., 1999), K-feldspar

(Wartho et al., 1999) and work is in progress on

estimates for plagioclase and leucite (Wartho and

Kelley, unpublished data). The current best estimates

are shown in Table 1 but some of the minerals most

commonly analysed in K–Ar dating, particularly

hydrous minerals such as muscovite, biotite and

amphibole, have yet to be characterised.

A strong compositional control of argon solubility

in minerals is unsurprising, and is comparable with

the strong compositional control on argon diffusion

(Dahl, 1996a,b). If argon solubility had been con-

trolled simply by the number of extended defects,

mineral composition would not have been an impor-

tant control on argon solubility. However, any dis-

cussion of excess argon in terms of argon solubility

and partition is inevitably a simplification of nature.

The increasing evidence for siting of noble gases in

lattice sites (e.g., Wood and Blundy, 2001; Chamorro

Table 1

Mineral Solubility (ppb bar� 1) Comments Source

Quartz 3.75 Maximum value, electron

microprobe determination

Roselieb et al., 1997

Clinopyroxene 0.09 Measured near detection limit Brooker et al., 1998

Olivine 1 Maximum value, (possible melt

inclusions in analysis)

Brooker et al., 1998

Phlogopite 1.8 Fluorine-rich mica Roselieb et al., 1999

K-feldspar 0.7 Low temperature ( < 750 jC) data Wartho et al., 1999

Plagioclase 0.20 Unpublished data Wartho and Kelley

Leucite Low T—60 Unpublished data Wartho and Kelley

High T—770

Fig. 4. Argon solubility ranges in hydrous fluid, melts and minerals.

Hydrous fluid solubility shows the range including high temperature

and high salinity data. Most melts fall in a relatively narrow range

but SiO2 has unusually high solubility for a melt of 1.8 ppm bar� 1.

Minerals show a wide range of solubility, probably correlated with

their ionic porosity, the mineral leucite has a high solubility as a

result of its unusual structure. Values for leucite are shown

separately from other minerals since both low and high temperature

forms show very high argon solubility.


et al., 2002) indicates the importance of vacancies and

thus intrinsic and extrinsic defects. Minerals contain-

ing potassium tend to have lattice vacancies suitable

for the large potassium ions, implying that argon

solubility in minerals is related to the vacancies and

may correlate with ionic porosity as it does in melts

and glasses (Carroll and Stolper, 1993). Moreover,

Dahl (1996a,b) has already demonstrated the power of

ionic porosity to predict differences in argon diffusiv-

ity between minerals. It is significant to the later

discussion of excess argon that the potassic minerals

(phlogopite, K-feldspar, and leucite) tend to exhibit

higher argon solubility than other minerals such as

plagioclase and clinopyroxene.

The ranges of argon solubility in hydrous fluids,

melts and minerals are compared in Fig. 4. Argon

solubility ranges over several orders of magnitude and

is highly incompatible in almost all mineral/melt or

mineral/fluid systems. Using the solubilities in

hydrous fluids and melts, and estimates of argon

solubility in minerals, it is possible to erect a model

of excess argon behaviour in simple natural systems.

3. Excess argon in open and closed systems

Argon partitioning between phases proves to be a

powerful method of understanding excess argon in

many systems. In particular, argon geochemistry

explains the presence of excess argon in fluid-poor

systems such as eclogites and blueschists (e.g., Scail-

let, 1996, 1998), and in fluid-rich systems such as

shear zones (e.g., Smith et al., 1994) or fluid-con-

trolled mineral deposits (e.g., Kendrick et al.,

2001a,b), in addition to explaining how most minerals

analysed for K–Ar and Ar–Ar dating are apparently

free from excess argon without the need for large fluid

fluxes along grain boundaries. Given that excess

argon has been detected in all types of crustal fluid,

and all but the youngest groundwater samples reflect

radiogenic argon input (e.g., Torgersen et al., 1989),

we might expect to find excess argon in many

minerals. This is of course exactly the situation in

other isotope geochronometers such as Rb–Sr, where

the initial 87Sr/86Sr ratio is a measure of the excess

radiogenic 87Sr in the system. The Sr isotope ratios of

the system components are measured as part of the

analysis, and it is not necessary to assume that Sr

exchanges with an infinite reservoir or escapes the

system, indeed closed system behaviour is common

(e.g.,, Jenkin et al., 1995; Jenkin, 1997 ).

The reason why K–Ar dating yields true ages in

the majority of cases is the highly incompatible

behaviour of the trace element argon (Fig. 4). Argon

strongly partitions from minerals into grain boundary

fluids in metamorphic rocks, or from crystals into

melts and melts into bubbles in magmatic systems,

leaving the minerals highly depleted in argon.

Although small amounts of excess argon may often

be present (unless the partition coefficients are infin-

ite), the concentrations must be so low as to be

swamped by in situ radiogenic argon. The highly

incompatible nature of the argon in solid/melt and

solid/fluid systems makes the fluids or melts effec-

tively ‘‘infinite reservoirs’’ for radiogenic argon in

many systems. The challenge of modelling argon

partitioning in natural systems, is to determine

whether routine K–Ar and Ar–Ar samples are on

the verge of detectable excess argon, or contain

negligible excess argon and those which do exhibit

excess argon are extreme end members. Until recently

such calculations have been impractical since the

measurement of argon solubility in minerals has been

unreliable. As a first attempt at understanding how the

system works, we can model the introduction of

excess argon into K-feldspar and use the results to

make important inferences about the susceptibility of

other minerals to excess argon.

Perhaps the most important concept introduced

when considering the behaviour of argon as a trace

element in a natural system, is that of open systems

and closed systems. Since argon is highly incompat-

ible, this question is central to understanding how

excess argon develops. An open system in this context

might be a shear zone through which there had been a

high fluid flux. If fluids flowing through the shear

zone were derived from ancient basement rocks and

contained high concentrations of radiogenic argon,

significant quantities of excess argon would partition

into minerals (e.g., Cumbest et al., 1994; Allen and

Stubbs, 1982). A closed system might be granulites

with very limited or transient fluid presence. In such a

fluid-poor closed system, transport of argon along the

grain boundaries might be as little as a few centi-

metres over millions of years (cf. Scaillet, 1996, 1998)

something which has also been demonstrated for


oxygen in HP and UHP rocks (Scaillet, 1996, 1998;

Philippot and Rumble, 2000). Above the closure tem-

perature, radiogenic argon produced in a closed sys-

tem would accumulate in the grain boundary network,

eventually building up to levels where dynamic equi-

librium would result in significant quantities residing

in the minerals. This scenario has been hypothesised

in several fluid-poor systems (e.g., Foland, 1979;

Scaillet, 1996; Kelley and Wartho, 2000; Baxter et

al., 2002). Although open and closed Ar systems are

very different, both can be successfully modelled by

considering argon as a trace element partitioning

between fluid and solid using the measured ranges

of argon solubility.

3.1. Excess argon in an open system

The key concern arising from argon mineral/fluid

partition is the extent to which the experimental

solubility data support an ‘infinite reservoir’ model

for argon loss in K–Ar geochronology. We can

investigate this using experimental and natural K-

feldspar data since this is the best experimentally

constrained example. To do this, we need to estimate

the relevant argon partition coefficients (KD) for K-

feldspar/saline fluid over a representative range of

conditions. Although there seems to be little variation

of argon solubility in K-feldspar with temperature

(Wartho et al., 1999), the Henry’s law coefficient of

water varies strongly with temperature (Figs. 2 and 3)

which means that KD will also vary with temperature.

In addition, the salting out effect lowers argon sol-

ubility in saline fluids (Figs. 2 and 3) and will also

cause variation in KD. Figs. 2 and 3 show the variation

of argon solubility in water with temperature (Cro-

vetto et al., 1982), at sufficiently high temperatures to

reach the range of closure for K-feldspar (around 210

jC for a 1-Am grains, and 260 jC for 10-Am grains

cooling at 10 jC/Ma). Argon solubility in saline

waters is significantly lower and has been measured

precisely only at lower temperatures (Smith and

Kennedy, 1983), but argon solubility for high temper-

atures can be estimated. Extrapolating the salinity data

in proportion with the high temperature solubility data

in Fig. 3, indicates that argon solubility in pure and

saline grain boundary fluids up to 300 jC lies in the

range 25–100 ppm bar � 1 atm � 1. If the solubility of

argon in K-feldspar is taken to be 0.7 ppb bar � 1, KD

for the K-feldspar grain boundary fluid system lies in

the range 7� 10� 6–3� 10 � 5. Given the lack of

variation in argon solubility in minerals (e.g., clino-

pyroxene, Chamorro et al., 2002; Kelley and Wartho,

unpublished data), the positive gradient of water

solubility vs. temperature (Fig. 3) means that for any

given system, KD between K-feldspar and the grain

boundary fluid varies antithetically with temperature.

An interesting feature of this temperature-con-

trolled KD variation is that in a cooling metamorphic

system, even in an open system, the amount of excess

argon introduced into K-feldspar will increase as the

temperature falls because argon solubility in the saline

grain boundary fluid decreases with consequent

increase of the partition coefficient. This effect is

likely to be universal for minerals since like melts,

they do not exhibit strong variation in solubility with

temperature. The common observation of late influxes

of excess argon into mineral grain boundaries (e.g.,

McDougall and Harrison, 1981; Pickles et al., 1997)

may not in fact reflect varying fluid composition, but

instead result from KD increase with cooling. A

similar observation can be made for the effect of

salinity variation in grain boundary fluid although it

may be less evident in natural samples. For example,

consider a saline grain boundary fluid containing

excess argon in equilibrium with excess argon in K-

feldspar. The introduction of more dilute surface water

would decrease the overall salinity, increase argon

solubility in the fluid and excess argon in the mineral

would diffuse into the fluid. Thus, excess argon can

be introduced or removed from minerals without a

large-scale flux of fluids through the Earth’s crust but

simply by temperature and salinity variation.

Returning to the question of the infinite reservoir,

Fig. 5 illustrates concentrations of excess argon in K-

feldspar (expressed as the increase in age they would

cause in K-feldspar) in equilibrium with argon in the

corresponding grain boundary fluid. The shaded area

indicates concentrations of (atmospheric) argon found

in near surface ground waters and some deeper waters

(Smith and Kennedy, 1983; Torgersen et al., 1989) for

comparison with the levels of excess argon in deeper

fluids. More extreme concentrations of excess argon

are found in hydrothermal fluids and in fluid inclu-

sions (circa 0.86 ppm, Kelley et al., 1986), micas

(Foland, 1979), and within K-feldspars (Turner and

Wang, 1992; Burgess et al., 1992; Harrison et al.,


1994). Harrison et al. (1994) calculated radiogenic

argon concentrations in fluid inclusions ranging from

0.08 ppm as high as 22 ppm based on an assumption

of fluid salinity (2%), although they stated that salinity

might be a factor of 5 higher (10%) indicating argon

concentrations perhaps as high as 110 ppm. Such

concentrations of excess argon in grain boundary

fluids would increase the apparent ages of K-feldspar

samples by 0.001 to 0.003 Ma at the lower end of the

observed natural Ar concentrations, to as high as 0.2

to 0.6 Ma at the upper end, probably undetectable in

all but the youngest samples. In the most extreme

case, with 110 ppm radiogenic argon in the fluid,

excess argon ages in the K-feldspar might amount to

age increases of 1.5 to 4 Ma. The highest levels of

radiogenic argon in K-feldspar-hosted fluid inclusions

were found only in older samples with protracted

histories (Harrison et al., 1994) and excess concen-

trations of 0.2 to 0.6 Ma would probably go unde-

tected in such samples. Thus, in the huge majority of

cases, the concept of an infinite reservoir works well

for K-feldspar in natural open systems. Although there

is a trivial amount of excess argon in all K-feldspars, it

is probably 1–2 orders of magnitude below detection

limits in most samples. This observation corroborates

the many measurements showing that excess argon is

uncommon in K-feldspar lattice, much of the excess

argon that is detected in K-feldspar is confined to fluid

inclusions. Another observation that can be derived

from Fig. 5 is the reason why K-feldspars generally

contain less atmospheric argon than hydrous minerals.

Using Fig. 5 to illustrate atmospheric argon rather

than excess argon indicates that a 100-Ma K-feldspar

would incorporate only 0.01% to 0.03% atmospheric

argon into the lattice from saline ground waters; thus,

even the small amounts of argon seen in K-feldspar

measurements are probably dominated by argon in

fluid inclusions and modern surface absorbed argon.

The model confirms how robust the K–Ar system

is when applied to K-feldspar, but what about other

potassium bearing minerals? In particular, how do

minerals with higher mineral/fluid partition coeffi-

cients behave in an open system and how does this

affect the infinite reservoir model in these cases?

Fig. 5. Excess argon ages in K-feldspar vs. argon concentration in a hypothetical infinite fluid. The shaded areas indicate likely range of

concentrations in natural fluids from groundwater and fluid inclusion studies. Most examples being above 0.1 ppm and below 10 ppm. The solid

lines indicate the likely upper and lower age limits of excess argon in K-feldspar using mineral/water partition coefficients of 7� 10� 6 to

3� 10� 5.The partition coefficients are so low that excess argon in K-feldspar will nearly always be below detection limits. The dashed lines

indicate partition for a mineral with higher mineral solubility, and thus partition coefficients of 1�10� 4 or 1�10� 3 such as biotite.


Biotite, in particular, has been shown to yield ages

over 100 Ma older than the expected age (e.g.,,

Brewer et al., 1969; Smith et al., 1994) and several

studies describe excess ages in biotite but not in co-

existing muscovite. Roddick et al. (1980) argued that

this reflected greater solubility of argon in biotite, an

observation corroborated by strong crystal chemical

(ionic porosity) arguments that argon solubility in

biotite should be higher than muscovite (Dahl,

1996b). Dashed lines in Fig. 5 illustrate the behaviour

of a mineral with a partition coefficients of 1�10� 4

and 1�10� 3, and similar radiogenic argon concen-

trations in grain boundary fluids to those used in the

K-feldspar model. The mineral would commonly

yield ages of the order of 1 Ma older than the true

closure age but under extreme conditions might yield

excess argon ages over 100 Ma older than the true

closure age (compared with 4 Ma for K-feldspar in the

same fluid). Although the highest partition coefficient,

1�10� 3 is still a factor of three smaller than that

measured by Onstott et al. (1991) for biotite, natural

studies have measured excess argon in metamorphic

biotite yielding ages over 100 million years older than

coexisting minerals (e.g., Brewer, 1969; Smith et al.,

1994), indicating that 1�10� 3 may be a reasonable

order of magnitude estimate. The extrapolated solu-

bility of argon in biotite is thus one or two orders of

magnitude higher than K-feldspar (Table 1), and

higher than the maximum value for quartz measured

by Roselieb et al. (1997), though recent work has

suggested that Ar solubility in quartz may be higher

(Watson and Cherniak, pers. comm.). The apparently

wide range of argon solubility in minerals is a

reasonable explanation of variations in excess argon

in fluid-rich environments where fluids are derived

from basement rocks, such as orogenic thrust belts

(e.g., Brewer et al., 1969; Kelley, 1988; Smith et al.,

1994; Reddy et al., 1997; Vance et al., 1998). The

widespread influx of fluids in such regimes is illus-

trated by the occurrence of similar concentrations of

excess argon in minerals over broad areas (e.g.,

Brewer et al., 1969; Smith et al., 1994).

3.2. Excess argon in a closed system

The development of excess argon in closed sys-

tems is a more recent discovery than fluid-borne

excess argon. Indeed, the only confirmed case for

many years was in fluid-poor rocks of the Arden

pluton, in the granulite-facies Wilmington Complex

in the Appalachian mountains (Foland, 1979). Rb–Sr

ages for biotite in these rocks were ca. 365 Ma

whereas K–Ar ages for biotites ranged from 365 to

590 Ma and correlated with whole rock potassium

content (Fig. 6A and B). The Arden pluton was

emplaced around 502 Ma, and remained above the

closure temperature of both biotite and K-feldspar

until 365 Ma. In fact, the rocks generally contain less

than 3% biotite, the variation in whole rock potassium

content is the result of variation in K-feldspar contents

(from near zero up to 41%). Biotites record K–Ar

ages older than the intrusive age and therefore contain

extraneous argon, but is it excess argon or inherited

Fig. 6. Excess argon in the Arden pluton (after Foland, 1979). (A)

K–Ar ages for biotite vs. whole rock potassium content, showing a

strong correlation. For comparison the Rb/Sr cooling age for biotite

(356 Ma) is shown. (B) Excess argon concentrations in biotite

pluton shown in ppb vs. whole rock potassium content.


argon? The extraneous argon could be described as

inherited or excess argon in this case, since the argon

is not retained within the biotite where it is observed

but is sourced within the rock. However, it will be

discussed as excess argon for the purposes of this

work. The strong correlation between excess argon

and whole rock potassium indicates a source within

the rock for the argon since it is heterogeneous on a

scale of as little as 10 m. However, the amount of

excess argon in biotite is always less than the amount

which would have been generated in the rock between

502 and 365 Ma. Even in the sample containing the

most excess argon (Fig. 6B), the biotite contains less

than 11% of the 40Ar generated between intrusion and

cooling. In fact, the simplest explanation of the excess

argon is that the incorporation of excess argon into

biotite reflects the relative solubility of argon in K-

feldspar>biotite>grain boundary fluid, even though

this may have been only a film of water covering

the grains. This sequence of relative solubility also

reflects experimental values (see previous section and

Table 1) and those found in UHP rocks (Arnaud and

Kelley, 1995). Although some excess argon in the

Arden pluton may have resided in the K-feldspar and

subsequently been lost (Foland, 1979), the correlation

between whole rock potassium and excess argon in

biotite means that the other ‘‘reservoir’’, most likely

the grain boundary network, was finite and contained

high concentrations of excess argon.

Scaillet (1998) reviewed the behaviour of excess

argon in high pressure (HP) and ultra-high pressure

(UHP) metamorphism and modelled the development

of excess argon in terms of a limited fluid phase.

Other work has shown that fluids in such rocks are

low volume, often transient, and they travel remark-

ably short distances in these environments (Philippot

and Rumble, 2000). This improved understanding of

fluid behaviour has been matched by an increasing

understanding of the behaviour of argon in such rocks

which can yield extreme ages reflecting the lack of

argon transport (Arnaud and Kelley, 1995; Scaillet,

1996). The form of fluid in the grain boundary net-

work of granulites, HP or UHP rocks is subject to

debate, and may only amount to an absorbed OH �

layer or even a CO2-rich fluid. Furthermore, any

fluids which were present may have been isolated in

pores or as grain edge tubules (Holness, 1997) and are

likely to have existed only transiently. However, if

such systems approach local equilibrium, as they

might at high temperature, they should be amenable

to a similar model to that utilised in the previous

section. This is a test of the limits of the ‘‘infinite

reservoir’’ model for K–Ar and Ar–Ar ages, applied

to fluid-poor systems. In a closed system, the control-

ling factors on the distribution of excess argon

between K-feldspar and fluid are: temperature, fluid

salinity, the volume fraction of fluid, and potassium

content or K-feldspar content of the whole rock. In

fluid-rich systems such as porous sandstones in a

basin environment, porosity might reach several per-

cent, but in metamorphic rocks, porosity decreases

with increasing grade and compaction to as little as

0.01% (10 � 4 volume fraction) in dry systems (Hol-

ness, 1997).

In this closed system model, zero radiogenic and

excess argon concentration is initially assumed in both

fluid and minerals. Although in natural systems there

is likely to have been some atmospheric argon, the

result will be equal partitioning of both 40Ar and 36Ar

into the mineral so we need only consider radiogenic40Ar for the moment. Argon behaviour in this simple

system has been modelled using the same range of K-

feldspar/fluid partition coefficients (7� 10 � 6 to

3� 10 � 5) to account for salinity variations. The

model has been run for rocks with 1% to 100% K-

feldspar, and argon is assumed to be even more

incompatible in mineral phases making up the rest

of the rock, making this a worst case scenario, con-

trasting with the Arden pluton example where biotite

was present. In this closed system, excess argon builds

up as radiogenic argon is produced by decay in K-

feldspar. In this case, the ‘‘reservoir’’ is finite but

takes time to fill with radiogenic argon to the level

where we would be able to detect excess in K-

feldspar. The fractional age excess has been calculated

for fluid filled porosity of 1%, 0.1% and 0.01% (Fig.

7). In samples with lowest porosity, the fractional

excess argon increases, showing that closed system

excess argon is more likely to develop in a fluid-poor

system. Argon is highly incompatible and thus with a

fluid filled porosity of 1%, even in a rock composed

of 100% K-feldspar, the fractional excess argon in K-

feldspar only reaches 0.07% to 0.26% of the duration

of the closed system. In other words, a 100-Ma old K-

feldspar suddenly exhumed from a closed system

would exhibit ages 0.07 to 0.26 Ma older than the


true age, probably undetectable. If the fluid filled

porosity is 0.1%, excess argon in the same K-feldspar

would yield ages from 0.7 to 3 Ma in a 100% K-

feldspar rock, but only 0.03 to 0.8 Ma in rocks within

the normal range of 5–30% K-feldspar in common

crustal rocks. In all probability, this would still be

below detection levels in most natural systems. Only

in the most fluid-poor systems with 0.01% porosity,

does the system start to exhibit detectable excess

argon with 2 to 8 Ma excess argon in a rock with

30% of a 100-Ma old K-feldspar, and even in this case

only the most K-feldspar-rich rocks containing very

saline fluids will produce detectable excess argon. K-

feldspar ages measured in eclogite terrains which

exhibit closed system excess argon in phengite

(Arnaud and Kelley, 1995), sometimes reveal high

temperature excess argon, although this might also

result from plagioclase or solid inclusions outgassing

during the cycle heating experiment (Arnaud and

Kelley, 1995; Boven et al., 2001). Furthermore, in

this model it was assumed that the other minerals in

the rock exhibited lower argon solubility than K-

feldspar but if another mineral is present with a higher

partition coefficient, such as biotite, excess argon

concentrations in K-feldspar quickly drop below

detection levels even in the most fluid-poor terrains

(e.g., Foland, 1979). It is not clear how far this model

can be extended into the most fluid-poor rocks since

the grain boundary fluid phase in dry systems such as

eclogites and granulites may be as little as a layer of

OH � molecules at the grain boundaries. However,

the model serves to demonstrate how a closed system

can explain phenomena observed in granulites, HP

and UHP terrains, particularly the occurrence of

excess argon in phengite when it is the predominant

potassium bearing mineral in the rock (e.g., Scaillet,

1998).

In several UHP terrains, it has been noted that

excess argon is more prevalent in rocks with pro-

tracted histories or old protoliths (Li et al., 1994;

Arnaud and Kelley, 1995; Inger et al., 1996; Scaillet,

1996; Sherlock et al., 1999; Giorgis et al., 2000).

Unlike the model above, such systems start with an

initial concentration of excess argon related to the age

Fig. 7. Fractional argon age increases in K-feldspar in a closed system with variable K-feldspar content. The solid, long dash and short dashed

lines indicate the age excess using K-feldspar/water partition coefficients of 7� 10� 6 to 3� 10� 5 and 1%, 0.1% and 0.01% free fluid in the

system. For systems containing 5% to 50% K-feldspar, the age excesses are only detectable in K-feldspar-rich systems with 0.01% hydrous

fluid. In systems with more fluid or containing other minerals with higher argon solubility, K-feldspar will exhibit even lower excess argon

concentrations.


and potassium content of the protolith. Although

much of this will partition into the limited fluid phase,

excess argon is more likely in rocks with older

protoliths. Another feature of a closed system is that

potassium-rich rocks will contain greater concentra-

tions of excess argon than potassium-poor rocks, even

when only one potassium bearing mineral phase is

present (Sherlock and Kelley, 2002). Ironically, in

fluid-poor high-pressure rocks it would be more

appropriate to measure ages from unpromising meta-

basalts or meta-sandstones rather than mica-rich

schists or K-feldspar-rich rocks. In fact meta-granites

containing both K-feldspar and mica are most likely to

contain excess in mica as shown by Foland (1979) and

Arnaud and Kelley (1995) and discussed by Scaillet

(1998).

Finally, what happens when the hydrous fluid in

the grain boundary network is replaced by a very thin

layer of melt in even higher temperature rocks? This

extreme example of argon behaviour in a closed

system may explain the ubiquitous presence of excess

argon in large phlogopite grains in kimberlite xeno-

liths (Phillips and Onstott, 1988; Phillips, 1991;

Kelley and Wartho, 2000). Phlogopite retains Ar in

most rocks at temperatures below 400 jC, where

diffusion is sufficiently slow in the mineral lattice.

However, the ages yielded by large phlogopites from

mantle and lower crustal xenoliths are commonly

older than the kimberlite eruption age, a phenomenon

which has been interpreted as the incorporation of

excess radiogenic Ar from a deep fluid source (Lover-

ing and Richards, 1964; Phillips and Onstott, 1988;

Phillips, 1991). However, Kelley and Wartho (2000)

showed that the Ar–Ar ages from the cores of large

phlogopite grains xenoliths brought to the surface by

kimberlites and diatremes were not random but cor-

responded with known events in the source region.

This result implies that phlogopites retained radio-

genic argon while remaining many hundreds of

degrees above their closure temperatures for extended

periods of time. Since there is no evidence that argon

diffusion in phlogopite could vary sufficiently to

cause such an effect, the only mechanism seems to

be a lack of Ar transfer to other phases or into the

grain boundary network. The rock acts as a small-

scale closed system, in other words the ‘infinite

reservoir’ criteria, assumed in most K–Ar and Ar–

Ar dating, is invalid in this rock. As in the UHP rocks,

xenoliths from the lower crust or lithospheric mantle

have extremely low free fluid concentrations. How-

ever, in this case the grain boundary space may be

filled not by saline fluids but by extremely thin melt

layers (Drury and FitzGerald, 1996; Wirth, 1996). In

this scenario, argon will partition into phlogopite in

preference to other more tightly packed mineral latti-

ces such as olivine, garnet or clinopyroxene. Even

though the rock seems a complex system, it can

actually be modelled as only two components, phlo-

gopite and grain boundary melt. An estimate of the

phlogopite/melt partition coefficient for argon in this

system can be obtained using the solubility of argon in

phlogopite of ca 1.8 ppb bar � 1 (Roselieb et al.,

1999). The grain boundary spaces into which Ar

would partition, are probably a few nanometres wide

representing around 0.0002 volume fraction of the

rock (Drury and FitzGerald, 1996; Wirth, 1996). The

best estimate of Ar solubility in the melt based on

studies at lower pressures is 20–100 ppb/bar (Carroll

and Stolper, 1993). Given these parameters, in a

closed system, 0.2–1% of the radiogenic argon would

partition into the grain boundary network at equili-

brium. Not only do the phlogopite grains retain radio-

genic argon in this closed system but they also retain

argon quantitatively. It is not clear whether all xen-

olith phlogopite core ages are meaningful, since fluids

in the upper mantle may contain significant quantities

of excess argon and thus transient fluids could intro-

duce excess argon into phlogopites. However, in the

absence of such fluids, the system is closed for

radiogenic argon.

4. Excess argon in fluid and melt inclusions

Fluid inclusions and melt inclusions are often

important sources of excess argon in minerals ana-

lysed for K–Ar and Ar–Ar dating, particularly in low

potassium minerals such as amphiboles and plagio-

clase. Mineral inclusions are excluded here though

these contribute in rare cases where mineral inclusions

retain older ages (e.g., Kelley et al., 1997), or prefer-

entially take up excess argon (e.g., Sisson and Onstott,

1986). The simple reason for the importance of fluid

and melt inclusions is illustrated by Fig. 4. Melt

inclusions in equilibrium with a magma containing

excess argon will contain circa. 100 times more argon


than the mineral lattice. Fluid inclusions in equili-

brium with a grain boundary fluid will contain as

much as 10,000 times the excess argon concentration

(by weight) of the mineral lattice. Fluid and melt

inclusions provide some of the most intractable ana-

lytical problems in K–Ar and Ar–Ar dating, but the

distribution of excess argon between inclusion and

mineral lattice conforms to the same simple rules as

those in open systems, described above.

4.1. Excess argon in fluid inclusions

Fluid inclusions have been known to be a source of

excess argon since the early days of K–Ar dating

(Rama et al., 1965) and their contents have been

quantified in quartz (Kelley et al., 1986), amphibole

(Cumbest et al., 1994) and K-feldspar (Burgess et al.,

1992; Turner and Wang, 1992; Harrison et al., 1994).

Fluid inclusions have significantly lower density than

their host minerals, for example, 10 100-Am fluid

inclusions in a 1-mm3 of K-feldspar represent circa.

0.5% by volume but only 0.2% by weight. However,

the relatively high solubility of argon in hydrous

fluids means that they can be very significant source

of excess argon to K–Ar dating even at the 0.5%

inclusion level.

In vacuo crushing has proved effective in assess-

ing excess argon in fluid inclusions, and has led to a

better understanding of the relationship between

argon and chlorine, using the Ar–Ar system. The

irradiation process creates 38Ar from 37Cl (Brereton,

1970), and thus facilitates not only direct measure-

ment of the argon to chlorine ratio, but also, given

measurements of fluid inclusion salinity, concentra-

tion of argon within the inclusion fluids. Kelley et al.

(1986) used a combination of in vacuo crushing and

stepped heating to quantify atmospheric, radiogenic

and excess components in quartz hosted fluid inclu-

sions formed from hydrothermal fluids, showing that

many samples had a roughly constant 40Ar*/Cl ratio

reflecting a similarly homogeneous salinity. Further

experiments with quartz, cherts and K-feldspars

(Turner and Bannon, 1992; Burgess et al., 1992;

Turner et al., 1993; Turner and Wang, 1994, Harrison

et al., 1994; Kendrick et al., 2001a,b) demonstrated

the ubiquity of the Ar/Cl relationship, although it is

only valid where fluid inclusions exhibit a small

variation in salinity.

Argon release by stepped heating has proved to be

a very effective technique for alleviating the problem

of excess argon in fluid inclusions, particularly in K-

feldspar (Harrison et al., 1994). The technique is

particularly effective in opening the larger inclusions

(greater than 5–10 Am) but decrepitation temperature

is a function of inclusion size (Shepherd et al., 1985),

such that larger inclusions tend to decrepitate at low

temperatures, and smaller inclusions decrepitate at

higher temperatures up to around 600 jC. Above thistemperature, the remaining inclusions are often less

than 1-Am diameter and their internal pressure never

reaches a level sufficient to cause decrepitation. The

sequence of increasing fluid inclusion decrepitation

temperatures was utilised by Harrison et al. (1994), to

reduce the effect of fluid inclusion interference in the

lower temperature argon release from K-feldspars

using repeated temperature steps and the correlation

of excess argon with chlorine to further correct the

repeat steps.

The excess argon contents of fluid inclusions

represent a range from circa 0.08 to 22 ppm, and

have varying effects upon the final age of the rock or

mineral, dependent upon the volume fraction of

inclusions and the potassium content of the host.

The consequences of fluid inclusions to K–Ar dating

are very variable (Fig. 8) but in almost all cases, the

effects are detectable. Only in the highest potassium

minerals, such as K-feldspar, with low excess argon

concentrations, would the effects be minimal. At the

lower end of excess argon concentrations in the fluid,

circa. 0.1 ppm, and 1% fluid inclusion presence in the

mineral, K-feldspar would exhibit age excesses of

around 0.35 Ma, whereas an amphibole (1% K) would

exhibit an age excess of 5 Ma. With fluids containing

10-ppm excess argon, not an unrealistic concentration

(Kelley et al., 1986; Harrison et al., 1994; Cumbest et

al., 1994), K-feldspars would exhibit excess ages of

35 Ma and amphiboles (1% K) would exhibit excess

ages of 500 Ma. The common observation that

amphiboles in orogenic belts yield anomalously old

K–Ar ages is probably often caused by fluid inclu-

sions. Cumbest et al. (1994) measured the effects of

fluid inclusions in amphiboles with a mean potassium

content of 0.59%. Stepped heating release patterns

yielded near plateau age release yielding ages of 70–

110 Ma older than the true cooling ages. They also

noted that inclusions did not all decrepitate during a


500 jC bake, leading to a ‘plateau-like’ release pattern

rather than a U- or saddle-shaped age pattern. Despite

anomalously high ages (500–700 Ma) in the first step,

the generally small size of the fluid inclusions, some

less than 1 Am, meant that they did not release argon

until the amphibole started to break down resulting in

plateau-like release. The amphiboles contained a het-

erogeneous population of fluid inclusions, of three

generations, each with an individual excess argon

signature. Cumbest et al. (1994) were able to separate

the signals from the three populations of fluid inclu-

sions showing that some low salinity inclusions were

rich in atmospheric argon whereas extremely saline

and excess argon-rich fluid inclusions contained fluids

derived from Proterozoic rocks in the footwall of a

shear zone.

4.2. Excess argon in melt inclusions

The importance of melt inclusions in volcanic

systems is less well recognised but well illustrated

by a detailed study of excess argon in recently erupted

(effectively zero age) anorthoclase on Mount Erebus

(Esser et al., 1997). While melt inclusions are gen-

erally easier to detect by observation than fluid

inclusions, and have relatively lower concentrations

of excess argon, they cannot easily be separated from

the mineral signal by stepped heating. Esser et al.

(1997) found that melt inclusions on the grain surfaces

released argon predominantly at temperatures of 900–

1000 jC, whereas buried melt inclusions released

their argon only when the anorthoclase started to

break down at around 1100 jC. Thus, minerals which

release argon at high temperature, such as K-feldspar,

plagioclase and amphibole, release much of their

inclusion derived excess argon synchronously with

the radiogenic argon release from the mineral lattice,

yielding anomalously high plateau ages. The effect of

melt inclusions is far less than fluid inclusions but can

be more important when dating young volcanic sam-

ples (e.g., Renne et al., 1997). For example, the zero

age Mount Erebus anorthoclase exhibited ages of 48,

179 and 640 ka in samples with f 1%, f 10% andf 30% melt inclusions (Esser et al., 1997). Fig. 9

illustrates the effect of 1% to 50% melt inclusions in

samples where the melt contains 0.1, 1 and 10 ppb

Fig. 8. Age excess vs. proportion of fluid inclusions (%) in minerals with variable potassium content. Solid lines indicate the excess argon age

increase in a mineral with 1%, 5% and 15% K and fluid inclusions containing 0.1-ppm excess argon. The short dashed lines indicate excess ages

in a mineral with 1%, 5% and 15% K and fluid inclusions containing 10-ppm excess argon.


excess argon. The three Mount Erebus samples lie

roughly on a line of single excess argon composition,

the slight disparity in the sample containing least

inclusions is probably the result of rounding of the

inclusion contents. Most minerals used to date very

young events, e.g., sanidine, contain more potassium

than the Mount Erebus anorthoclase and are thus less

affected. A sanidine containing melt inclusions with

similar excess argon concentrations to Mt. Erebus

would exhibit excess ages of only 1, 10 and 30 ka

(note that Renne et al., 1997 detected very small

amounts of excess argon in sanidine from Vesuvius).

If the Mount Erebus anorthoclase had been 1 Ma old,

age excesses would have been insignificant in the 1%

inclusion sample and barely detectable in the 10%

melt inclusion sample.

5. Conclusions

The occurrence of excess argon in K–Ar and Ar–

Ar geochronology can be understood by considering

argon as an incompatible trace element, exhibiting

mineral/fluid and mineral/melt partition coefficients

ranging from 0.01 to as low as 7� 10 � 6. The

extremely incompatible nature of argon is the main

reason why excess argon is a relatively uncommon

phenomenon. Argon tends to partition into melt or

hydrous fluid in the grain boundary network and

even in low porosity rocks this reservoir dominates

the system. This view of excess argon supports the

long-held concept of argon escaping from minerals

into an ‘infinite reservoir’. Excess argon results

when argon concentrations in this melt or fluid

reservoir are sufficiently high that dynamic equili-

brium leads to significant argon partitioning into mi-

nerals.

By discussing argon as trace element with known

solubility in hydrous fluids, melts and minerals, it is

possible to quantify excess argon in natural systems as

a series of mineral/fluid partition coefficients. Models

for K-feldspar in both open and closed systems

confirm that it will only exhibit excess argon in the

most extreme circumstances whereas minerals with

higher partition coefficient such as biotite may more

commonly yield anomalously old ages. The models

show that the ‘infinite reservoir’ model, which allows

argon to escape from minerals into grain boundary

Fig. 9. Age excess vs. proportion of melt inclusions (%) in minerals with variable potassium content. Solid lines indicate the excess argon age

increase for a mineral with 1%, 5% and 15% K and a fluid containing 0.1-ppb excess argon. Long dashed lines indicate excess ages for a

mineral with 1%, 5% and 15% K and a fluid containing 1-ppb excess argon. Three data points are plotted for anorthoclase containing f 1%,

f 10% and f 30% melt inclusions from Mount Erebus (Esser et al., 1997).


fluids and be lost from the system, is robust and does

not lead to significant errors in most cases. The

models also predict several features observed in

nature, in particular cooling in metamorphic terrains

is paralleled by increasing partition mineral/fluid

coefficients. Thus, the commonly observed late intro-

duction of excess argon into mineral grain boundaries

may be caused by cooling in a closed system and not

by an influx of excess argon-rich fluids.

In fluid-poor closed systems such as granulites, HP

and UHP rocks, excess argon builds up in response to

the potassium content of the rock. Although K-feld-

spar only exhibits excess argon in extreme circum-

stances, hydrous minerals such as phengite in dry

UHP rocks and phlogopite in mantle xenoliths are

more susceptible. Although phengites in UHP rocks

or biotites in granulites may retain small proportion of

the total radiogenic argon generated in the rock,

phlogopites in the lower crust and lithospheric mantle

may retain argon quantitatively, with the ‘excess’

argon reflecting storage times in the mantle since

the melt at the grain boundaries has much lower argon

solubility than hydrous fluids.

Excess argon in fluid and melt inclusions contrib-

utes a significant portion of excess argon in low

potassium minerals such as amphiboles and plagio-

clase. Although melt inclusions cause smaller

excesses of argon than fluid inclusions, because of

the lower melt/crystal partition coefficient (KD), this is

a particular problem in young volcanics.

Acknowledgements

I am greatly indebted to Jo-Anne Wartho and Phil

Guise for encouraging me to write this article. The

manuscript was improved by discussions with Nicolas

Arnaud and Sarah Sherlock, and two thorough and

reviews from Stephane Scaillet and Jim Lee. [SK]

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Excess argon in K–Ar and Ar–Ar geochronology

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